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)029<0319:AWGITN>2.0.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)024<0439: JAAMIS>2.3.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
The assembly of West Gondwana—The view from the Rio de la Plata craton foreland domain, the Rio de la Plata craton. Framework, tectonic evolution and correlation with similar provinces of southwestern Africa, 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: Rio de Janeiro, Sociedade Brasileira de Geologia, p. 311–334. Basei, M.A.S., Frimmel, H.E., Nutman, A.P., Preciozzi, F., and Jacob, J., 2005, A connection between the Neoproterozoic Dom Feliciano (Brazil/Uruguay) and Gariep (Namibia/South Africa) orogenic belts—Evidence from a reconnaissance provenance study: Precambrian Research, v. 139, p. 195–221, doi: 10.1016/j.precamres.2005.06.005. Boger, S.D., and Miller, J.McL., 2004, Terminal suturing of Gondwana and the onset of the Ross–Delamerian orogeny: The cause and effect of an Early Cambrian reconfiguration of plate motions: Earth and Planetary Science Letters, v. 219, p. 35–48, doi: 10.1016/S0012-821X(03)00692-7. Boger, S.D., Carson, C.J., Fanning, C.M., Hergt, J.M., Wilson, C.J.L., and Woodhead, J.D., 2002, Pan-African intraplate deformation in the northern Prince Charles Mountains, east Antarctica: Earth and Planetary Science Letters, v. 195, p. 195–210. Brito Neves, B.B., Campos Neto, M.C., and Fuck, R.A., 1999, From Rodinia to Western Gondwana: An approach to the Brasiliano-Pan African cycle and orogenic collage: Episodes, v. 22, p. 155–166. Brueckner, H.K., Cunningham, D., Alkmin, F.F., and Marshak, S., 2000, Tectonic implications of Precambrian Sm-Nd dates from the southern São Francisco craton and adjacent Araçuaí and Ribeira belts, Brazil: Precambrian Research, v. 99, p. 255–269, doi: 10.1016/S0301-9268(99)00065-0. Caby, R., Arthaud, M.H., and Arcanjo, C.J., 1995, Lithostratigraphy and petrostructural characterization of supracrustal units in the Brasiliano belt of northeastern Brazil: Geodynamic implications: Journal of South American Earth Sciences, v. 8, p. 235–246, doi: 10.1016/0895-9811(95)00011-4. Campos Neto, M.C., and Figueiredo, M.C.H., 1995, The Rio Doce orogeny, southeastern Brasil: Journal of South American Earth Sciences, v. 8, p. 143–162. Chemale, F., Jr., 2000, Evolução Geológica do Escudo Sul-rio-grandense, in Holz, M., and De Ros, L.F., eds., Geologia do Rio Grande do Sul: Porto Alegre, Brazil, CIGO/Universidade Federal do Rio Grande do Sul, p. 13–52. Chemale, F., Jr., Hartmann, L.A., and da Silva, L.C., 1995, Stratigraphy and tectonism of the Brasiliano cycle in southern Brazil: Communications Geological Survey of Namibia, v. 10, p. 151–166. Cordani, U.G., D’Agrella-Filho, M.S., Brito-Neves, B.B., and Trindade, R.I.F., 2003, Tearing up Rodinia: The Neoproterozoic palaeogeography of South American cratonic fragments: Terra Nova, v. 15, p. 350–359, doi: 10.1046/j.1365-3121.2003.00506.x. Cosca, M.A., Shimron, A., and Caby, R., 1999, Late Precambrian metamorphism and cooling in the Arabian-Nubian Shield: Petrology and 40 Ar/39Ar geochronology of metamorphic rocks of the Elat area (southern Israel): Precambrian Research, v. 98, p. 107–127, doi: 10.1016/S03019268(99)00044-3. Costa, A.F.U., 1998, Geophysical modeling and testing of the structure of the lithotectonic associations in the Sul-Riograndense Shield, southern Brazil: Pesquisas, v. 25, no. 2, p. 75–76. Crawford, A.J., and Berry, R.F., 1992, Tectonic implications of Late Proterozoic–Early Paleozoic igneous rock associations in western Tasmania: Tectonophysics, v. 214, p. 37–56, doi: 10.1016/0040-1951(92)90189-D. Dalla Salda, L.H., Cingolani, C.A., and Varela, R., 1992, Early Paleozoic orogenic belt on the Andes in South America: Results of Laurentia-Gondwana collision? Geology, v. 20, p. 616–620. Dalziel, I.W.D., Mosher, S., and Gahagan, L., 2000, Laurentia-Kalahari collision and the assembly of Rodinia: Journal of Geology, v. 108, p. 499–513, doi: 10.1086/314418. da Silva, L.C., Hartmann, L.A., McNaughton, N.J., and Fletcher, I.R., 1999, SHRIMP U/Pb zircon dating of Neoproterozoic granitic magmatism and collision in the Pelotas Batholith, southernmost Brazil: International Geology Review, v. 41, p. 531–551. da Silva, L.C., Gresse, P.G., Scheepers, R., McNaughton, N.J., Hartmann, L.A., and Fletcher, I.R., 2000a, U-Pb SHRIMP and Sm-Nd age constraints on the timing and sources of the Pan-African Cape Granitic Suite, South Africa: Journal of African Earth Sciences, v. 30, p. 795–815, doi: 10.1016/S0899-5362(00)00053-1. da Silva, L.C., Hartmann, L.A., McNaughton, N.J., and Fletcher, I.R., 2000b, Zircon U/Pb SHRIMP dating of a Neoproterozoic overprint in Paleoproterozoic granitic-gneissic terranes, southern Brazil: American Mineralogist, v. 85, p. 649–668. da Silva, L.C., McNaughton, N.J., Armstrong, R., Hartmann, L.A., and Fletcher, I.R., 2005, The neoproterozoic Mantiqueira Province and its African
23
connections: A zircon-based U-Pb geochronologic subdivision for the Brasiliano/Pan-African systems of orogens: Precambrian Research, v. 136, p. 203–240, doi: 10.1016/j.precamres.2004.10.004. De Witt, M.T., Thiart, C., Doucouret, M., and Wilsher, W., 1999, Scent of a supercontinent: Gondwana’s ores as chemical tracers—tin tungsten and the Neoproterozoic Laurentia-Gondwana connection: Journal of African Earth Sciences, v. 28, p. 35–51, doi: 10.1016/S0899-5362(98)00085-2. Doig, R., Murphy, J.B., and Nance, R.D., 1993, Tectonic significance of the Late Proterozoic Economy River Gneiss, Cobequid Highlands, Avalon composite terrane, Nova Scotia: Canadian Journal of Earth Sciences, v. 30, p. 474–479. Dörr, W., Finger, F., Linnemann, U., and Zulauf, G. (eds.), 2004, The Avalonian-Cadomian belt and related peri-Gondwana terranes: International Journal of Earth Sciences, v. 93, p. 657–944. Dürr, S.B., and Dingeldey, D.P., 1996, The Kaoko belt (Namibia): Part of a late Neoproterozoic continental-scale strike-slip system: Geology, v. 24, p. 503– 506, doi: 10.1130/0091-7613(1996)024<0503:TKBNPO>2.3.CO;2. Egal, E., Guerrot, C., Le Goff, D., Thieblemont, D., and Chantraine, J., 1996, The Cadomian orogeny revisited in northern France, in Nance, R.D., and Thompson, M.D., eds., Avalonian and related peri-Gondwanan terranes of the circum-North Atlantic: Boulder, Colorado, Geological Society of America Special Paper 304, p. 281–318. Fernandes, L.A.D., and Koester, E., 1999, The Neoproterozoic Dorsal de Canguçu strike-slip shear zone: Its nature and role in the tectonic evolution of southern Brazil: Journal of African Earth Sciences, v. 29, p. 3–24, doi: 10.1016/S0899-5362(99)00076-7. Fernandes, L.A.D., Tommasi, A., and Porcher, C.C., 1992, Deformation patterns in the southern Brazilian branch of the Dom Feliciano belt: A reappraisal: Journal of South American Earth Sciences, v. 5, no. 1, p. 77–96, doi: 10.1016/0895-9811(92)90061-3. Fernandes, L.A.D., Menegat, R., Costa, A.F.U., Koester, E., Kramer, G., Tommasi, A., Porcher, C.C., Ramgrab, G.E., and Camozzato, E., 1995, Evolução tectônica do Cinturão Dom Feliciano no Escudo Sul-rio-grandense: Parte I uma contribuição a partir do registro geológico: Revista Brasileira Geosciências, v. 25, p. 351–374. Ferreira Filho, C.F., Kamo, S.L., Fuck, R.A., Krogh, T.E., and Naldrett, A.J., 1994, Zircon and rutile U-Pb geochronology of the Niquelandia layered mafic and ultramafic intrusion, Brazil: Constraints for the timing of magmatism and high grade metamorphism: Precambrian Research, v. 68, p. 241–255, doi: 10.1016/0301-9268(94)90032-9. Finger, F., Hanzl, P., Pin, C., von Quadt, A., and Steyrer, H.P., 2000, The Brunovistulian: Avalonian Precambrian sequence at the eastern end of the central European Variscides? in Franke, W., Haak, V., Oncken, O., and Tanner, D., eds., Orogenic processes: Quantification and modelling in the Variscan belt: Geological Society of London Special Publication 179, p. 103–112. Finney, S., Gleason, J., Gehrels, G., Peralta, S., and Acenõlaza, G., 2003, Early Gondwanan connection for the Argentine Precordillera terrane: Earth and Planetary Science Letters, v. 205, p. 349–359. Foden, J., Sandiford, M., Dougherty-Page, J., and Williams, I., 1999, Geochemisty and geochronology of the Rathjen gneiss: Implications for the early tectonic evolution of the Delamerian orogen: Australian Journal of Earth Sciences, v. 46, p. 377–389, doi: 10.1046/j.1440-0952.1999.00712.x. Fragoso-César, A.R.S., 1991, Tectônica de Placas no Ciclo Brasiliano: As orogenias dos Cinturões Dom Feliciano e Ribeira no Rio Grande do Sul [Ph.D. thesis]: São Paulo, University of São Paulo, 367 p. Frantz, J.C., and Botelho, N.F., 2000, Neoproterozoic granitic magmatism and evolution of the eastern Dom Feliciano belt in southernmost Brazil: A tectonic model: Gondwana Research, v. 3, p. 7–19, doi: 10.1016/S1342937X(05)70053-8. Frantz, J.C., and Nardi, L.V.S., 1992, O Magmatismo Granítico da região Oriental do Escudo Sul-rio-grandense. Uma Revisão: Pesquisas, v. 19, p. 183–189. Frantz, J.C., Botelho, N.F.Q., Pimentsl, M.M., Potrel, A., Koester, E., and Teixeira, R.S., 1999, Relações isotópicas Rb-Sr e Sm-Nd e idades do magmatsimo granítico Brasiliano da região leste do cinturão Dom Feliciano no Rio Grande do Sul: Evidêcias de retrabalhamento de crosta continental Paleoproterozóica: Revista Brasileira de Geociências, v. 29, p. 227–232. Frantz, J.C., Remus, M.V.D., and Hartmann, L.A., 2000, Geological units, ages and tectonic evolution of the Neoproterozoic Dom Feliciano belt, southernmost Brazil—A review: Revista Brasileiras de Geociências, v. 30, p. 55–57. Frimmel, H.E., and Fölling, P.G., 2004, Late Vendian closure of the Adamastor Ocean: Timing of tectonic inversion and syn-orogenic sedimentation in the Gariep basin: Gondwana Research, v. 7, p. 685–699, doi: 10.1016/ S1342-937X(05)71056-X.
24
Saalmann et al.
Frimmel, H.E., and Frank, W., 1998, Neoproterozoic tectono-thermal evolution of the Gariep belt and its basement: Namibia and South Africa: Precambrian Research, v. 90, p. 1–28. Frimmel, H.E., and Zartmann, R.E., 2000, The Richtersveld Igneous Complex: Beginning of Pan-African continental break-up in South Africa: Journal of South African Earth Sciences, v. 31, p. 22–23. Frimmel, H.E., Hartnady, C.J.H., and Koller, F., 1996, Geochemistry and tectonic setting of magmatic units in the Pan-African Gariep belt, Namibia: Chemical Geology, v. 130, p. 101–121, doi: 10.1016/0009-2541(95)00188-3. Gibson, G.M., and Wright, T.O., 1985, Importance of thrust faulting in the tectonic development of northern Victoria Land, Antarctica: Nature, v. 315, p. 480–483, doi: 10.1038/315480a0. Giese, U., and Buehn, B., 1994, Early Paleozoic rifting and bimodal volcanism in the Ossa-Morena Zone of south-west Spain: Geologische Rundschau, v. 83, p. 143–160, doi: 10.1007/BF00211898. Goodge, J.W., 1997, Latest Neoproterozoic basin inversion of the Beardmore Group, central Transantarctic Mountains, Antarctica: Tectonics, v. 16, p. 682–701, doi: 10.1029/97TC01417. Goodge, J.W., and Dallmeyer, R.D., 1992, 40Ar/39Ar mineral age constraints on the Paleozoic tectonothermal evolution of high-grade basement rocks within the Ross orogen, central Transantarctic Mountains: Journal of Geology, v. 100, p. 91–106. Goscombe, B., Hand, M., and Gray, D., 2003, Structure of the Kaoko belt, Namibia: Progressive evolution of a classic transpressional orogen: Journal of Structural Geology, v. 25, p. 1049–1081, doi: 10.1016/S0191-8141(02)00150-5. Grunow, A., Hanson, R., and Wilson, T., 1996, Were aspects of Pan-African deformation linked to Iapetus opening? Geology, v. 24, p. 1063–1066, doi: 10.1130/0091-7613(1996)024<1063:WAOPAD>2.3.CO;2. Guerrot, C., and Peucat, J.J., 1990, U-Pb geochronology of the Upper Proterozoic Cadomian orogeny in the northern Armorican Massif, France, in D’Lemos, R.S., Strachan, R.A., and Topley, C.G., eds., The Cadomian orogeny: Geological Society of London Special Publication 51, p. 13–26. Hartmann, L.A., 1998, Deepest exposed crust of Brazil—Geochemistry of Paleoproterozoic depleted Santa Maria Chico granulites: Gondwana Research, v. 1, p. 331–341, doi: 10.1016/S1342-937X(05)70849-2. Hartmann, L.A., and Delgado, I.M., 2001, Cratons and orogenic belts of the Brazilian shield and their contained gold deposits: Mineralium Deposita, v. 36, p. 207–217, doi: 10.1007/s001260100175. Hartmann, L.A., and Nardi, L.V.S., 1983, Contribuição à geologia da região oeste do Escudo Sul-Riograndense, in 1st Simpósio Sul-Brasileiro de Geologia: Porto Alegre, Brazil, Sociedade Brasileira de Geologia, p. 9–18. Hartmann, L.A., and Remus, M.V.D., 2000, Origem e Evolução das Rochas Ultramáficas do Rio Grande do Sul desde o Arqueano até o Cambriano, in Holz, M., and De Ros, L.F., eds., Geologia do Rio Grande do Sul: Porto Alegre, Brazil, CIGO/Universidade Federal do Rio Grande do Sul, p. 13–52. Hartmann, L.A., Nardi, L.V.S., Formoso, L.L., Remus, M.V.D., De Lima, E.F., and Mexias, A.S., 1999, Magmatism and metallogeny in the crustal evolution of Rio Grande do Sul Shield, Brazil: Pesquisas, v. 26, p. 45–63. Hartmann, L.A., Leite, J.A.D., da Silva, L.C., Remus, M.V.D., McNaughton, N.J., Groves, D.I., Fletcher, I.R., Santos, J.O.S., and Vasconcellos, M.A.Z., 2000a, Advances in SHRIMP geochronology and their impact on understanding the tectonic and metallogenic evolution of southern Brazil: Australian Journal of Earth Sciences, v. 47, p. 829–844, doi: 10.1046/ j.1440-0952.2000.00815.x. Hartmann, L.A., Porcher, C.C., and Remus, M.V.D., 2000b, Evolução das rochas metamórficas do Rio Grande do Sul, in Holz, M., and DeRos, L.F., eds., Geologia do Rio Grande do Sul: Porto Alegre, Brazil, CIGO/Universidade Federal do Rio Grande do Sul, p. 79–118. Hartmann, L.A., Philipp, R.P., Liu, D., Wang, L., Santos, J.O.S., and Vasconcellos, M.A.Z., 2004, Paleoproterozoic magmatic provenance of detrital zircons, Porongos Complex quartzites, Southern Brazilian Shield: International Geology Review, v. 46, p. 127–157. Hartnady, C., Joubert, P., and Stowe, C., 1985, Proterozoic crustal evolution in southwestern Africa: Episodes, v. 8, no. 4, p. 236–244. Heilbron, M., and Machado, N., 2003, Timing of terrane accretion in the Neopoterozoic-Eopaleozoic Ribeira orogen (SE Brazil): Precambrian Research, v. 125, p. 87–112, doi: 10.1016/S0301-9268(03)00082-2. Heilbron, M., Pedrosa-Soares, A., Neto, M., da Silva, L., Trouw, R., and Janasi, V., 2004, Brasiliano orogens in southeast and south Brazil, in Weinberg, R., Trouw, R., Fuck, R., and Hackspacher, P., eds., The 750–550 Ma Brasiliano event of South America: Journal of the Virtual Explorer, electronic edition, ISSN 1441–8142, v. 17, paper 4. http://www. virtualexplorer.com.au/journal/2004/17/.
Hoffmann, P.F., Hawkins, D.P., Isachsen, C.E., and Bowring, S.A., 1996, Precise U-Pb zircon ages for early Damaran magmatism in the Summas Mountains and Weltwitschia inlier, northern Damara belt, Namibia: Communications of the Geological Survey of Namibia, v. 11, p. 47–72. Hölzl, S., Hofmann, A.W., Todt, W., and Köhler, H., 1994, U-Pb geochronology of the Sri Lankan basement: Precambrian Research, v. 66, p. 123–149. Iacumin, M., Piccirillo, E.M., Girardi, V.A.V., Teixeira, W., Bellini, G., Echeveste, H., Fernandez, R., Pinese, J.P.P., and Ribot, A., 2001, Early Proterozoic calc-alkaline and Middle Proterozoic thoeliitic dyke swarms from central-eastern Argentina: Petrology, geochemistry, Sr-Nd isotopes and tectonic implications: Journal of Petrology, v. 42, p. 2109–2143, doi: 10.1093/petrology/42.11.2109. Issler, R.S., 1983, Evolução crustal da Faixa Arco-Fossa Tijucas e Faixa Magmática Pedras Grandes: Cráton Dom Feliciano, in 1st Simpósio SulBrasileiro de Geologia: Porto Alegre, Brazil, Sociedade Brasileira de Geologia, p. 19–36. Jacobs, J., Fanning, C.M., Henjes-Kunst, F., Olesch, M., and Paech, H., 1998, Continuation of the Mozambique belt into east Antarctica: Grenville-age metamorphism and polyphase Pan-African high-grade events in Central Dronning Maud Land: Journal of Geology, v. 106, p. 385–406. Jost, H., 1982, Condições do metamorfismo regional de uma parte da faixa de dobramentos de Tijucas no Rio Grande do Sul-RS: Acta Geologica Leopoldensia, v. 12, p. 3–32. Jost, H., and Bitencourt, M.F., 1980, Estratigrafia e tectônica de uma fração da Faixa de Dobramentos de Tijucas no Rio Grande do Sul: Acta Geologica Leopoldensia, v. 4, p. 27–60. Jost, H., and Hartmann, L.A., 1984, Provincia Mantiqueira—Setor Meridional, in Almeida, F.F.M., and Hasui, Y., eds., O Pré-Cambriano do Brasil: São Paulo, Edgard Blucher, p. 345–368. Jung, S., and Mezger, K., 2003, Petrology of basement-dominated terranes: I. Regional metamorphic T-t path from U-Pb monazite and Sm-Nd garnet geochronology (Central Damara orogen, Namibia): Chemical Geology, v. 198, p. 223–247, doi: 10.1016/S0009-2541(03)00037-8. Junges, S.L., Pimentel, M.M., and de Moraes, R., 2002, Nd isotopic study of the Neoproterozoic Mara Rosa arc, central Brazil: Implications for the evolution of the Brasília belt: Precambrian Research, v. 117, p. 101–118, doi: 10.1016/S0301-9268(02)00077-3. Kleinschmidt, G., and Tessensohn, F., 1987, Early Paleozoic westward directed subduction at the Pacific margin of Antarctica, in McKenzie, G.D., ed., Gondwana six: Structure, tectonics, and geophysics: Washington, D.C., American Geophysical Union, p. 89–105. Koester, E., Soliani, E., Jr., Fernandes, L.A.D., Kraemer, G., and Tommasi, A., 1997, Geocronologia Rb-Sr e K-Ar dos granitóides sintectônicos à Zona de Cisalhamento Transcorrente Dorsal de Canguçu na região de Encruzilhada do Sul: RS: Pesquisas, v. 24, p. 67–77. Koppe, J.C., and Hartmann, L.A., 1988, Geochemistry of the Bossoroca greenstone belt, southernmost Brazil: Geochimica Brasiliensis, v. 2, p. 167–174. Kröner, A., and Cordani, U., 2003, African and South American cratons were not part of the Rodinia supercontinent: Evidence from field relationships and geochronology: Tectonophysics, v. 375, p. 325–352, doi: 10.1016/ S0040-1951(03)00344-5. Kröner, A., Willner, A.P., Hegner, E., Jaeckel, P., and Nemchin, A., 2001, Single zircon ages, PT evolution and Nd isotopic systematics of highgrade gneisses in southern Malawi and their bearing on the evolution of the Mozambique belt in southern Africa: Precambrian Research, v. 109, p. 257–291, doi: 10.1016/S0301-9268(01)00150-4. Laux, J.H., Pimentel, M.M., Dantas, E.L., Armstrong, R., and Junges, S.L., 2005, Two neoproterozoic crustal accretion events in the Brasília belt, central Brazil: Journal of South American Earth Sciences, v. 18, p. 183–198, doi: 10.1016/j.jsames.2004.09.003. Leite, J.A.D., Hartmann, L.A., McNaughton, N.J., and Chemale, F., Jr., 1998, SHRIMP U/Pb zircon geochronology of Neoproterozoic juvenile and crustal-reworked terranes in southernmost Brazil: International Geology Review, v. 40, p. 688–705. Leite, J.A.D., Hartmann, L.A., Fernandes, L.A.D., McNaughton, N.J., Soliani, E.Ã., Jr., Koester, E., Santos, J.O.S., and Vasconcellos, M.A.Z., 2000, Zircon U-Pb SHRIMP dating of gneissic basement of the Dom Feliciano belt, southernmost Brazil: Journal of South American Earth Sciences, v. 13, p. 739–750, doi: 10.1016/S0895-9811(00)00058-4. Linnemann, U., Gehmlich, M., Tichomirowa, M., Buschmann, B., Nasdala, L., Jonas, P., Lützner, H., and Bombach, K., 2000, From Cadomian subduction to Early Palaeozoic rifting: The evolution of Saxo-Thuringia at the margin of Gondwana in the light of single zircon geochronology and basin development (central European Variscides, Germany), in
The assembly of West Gondwana—The view from the Rio de la Plata craton Franke, W., Haak, V., Oncken, O., and Tanner, D., eds., Orogenic processes: Quantification and modelling in the Variscan belt: Geological Society of London Special Publication 179, p. 131–153. Machado, N., Koppe, J.C., and Hartmann, L.A., 1990, A Late Proterozoic U-Pb age for the Bossoroca belt, Rio Grande do Sul, Brazil: Journal of South American Earth Sciences, v. 3, p. 87–90, doi: 10.1016/08959811(90)90021-R. Machado, N., Valladares, C., Heilbron, M., and Valeriano, C., 1996, U-Pb geochronology of the central Ribeira belt (Brazil) and implications for the evolution of the Brazilian orogeny: Precambrian Research, v. 79, p. 347– 361, doi: 10.1016/0301-9268(95)00103-4. Mantovani, M.S.M., Hawkesworth, C.J., and Basei, M.A.S., 1987, Nd and Pb isotope studies bearing on crustal evolution of southeastern Brazil: Revista Brasileira de Geosciências, v. 17, p. 263–268. Marques, J.C., Roisenberg, A., and Jost, H., 1996, Cromoespinélio como indicador petrogenético, Suíte Ultramáfica Cerro da Mineração: Congresso Brasileiro de Geologia 39, Salvador: Boletim de Resumos Expandidos Salvardor, Sociedade Brasileira de Geologia, v. 6, p. 221–223. Martins, V.T. de S., Teixeira, W., Noce, C.M., and Pedrosa-Soares, A.C., 2004, Sr and Nd characteristics of Brasiliano/Pan-African granitoid plutons of the Araçuaí orogen, southeastern Brazil: Tectonic implications: Gondwana Research, v. 7, p. 75–89, doi: 10.1016/S1342-937X(05)70307-5. Meert, J.G., 2003, A synopsis of events related to the assembly of eastern Gondwana: Tectonophysics, v. 362, p. 1–40, doi: 10.1016/S0040-1951(02)00629-7. Meert, J.G., and Van der Voo, R., 1997, The assembly of Gondwana 800–500 Ma: Journal of Geodynamics, v. 23, p. 223–235, doi: 10.1016/S02643707(96)00046-4. Meert, J.G., Van der Voo, R., and Ayub, S., 1995, Paleomagnetic investigation of the Neoproterozoic Gagwe lavas and Mbozi Complex, Tanzania, and the assembly of Gondwana: Precambrian Research, v. 74, p. 225–244, doi: 10.1016/0301-9268(95)00012-T. Miller, J.S., Santoch, M., Pressley, R.A., Clements, A.S., and Rogers, J.W., 1996, A Pan-African thermal event in southern India: Journal of Southeast Asian Earth Sciences, v. 14, p. 127–136, doi: 10.1016/S0743-9547(96)00051-7. Möller, A., Mezger, K., and Schenk, V., 2000, U-Pb dating of metamorphic minerals: Pan-African metamorphism and prolonged slow cooling of high pressure granulites in Tanzania, east Africa: Precambrian Research, v. 104, p. 123–146, doi: 10.1016/S0301-9268(00)00086-3. Münker, C., and Crawford, A.J., 2000, Cambrian arc evolution along the SE Gondwana active margin: A synthesis from Tasmania-New ZealandAustralia-Antarctica correlations: Tectonics, v. 19, p. 415–432, doi: 10.1029/2000TC900002. Murphy, J.B., Keppie, J.D., Dostal, J., and Nance, R.D., 1999, Neoproterozoic– early Paleozoic evolution of Avalonia, in Ramos, V.A., and Keppie, J.D., eds., Laurentia-Gondwana connections before Pangea: Boulder, Colorado, Geological Society of America Special Paper 336, p. 253–266. Myrow, P.M., Pope, M.C., Goodge, J.W., Fischer, W., and Palmer, A.R., 2002, Depositional history of pre-Devonian strata and timing of Ross orogenic tectonism in the central Transantarctic Mountains, Antarctica: Geological Society of America Bulletin, v. 114, p. 1070–1088. Nance, R.D., Murphy, J.B., Strachan, R.A., D’Lemos, R.S., and Taylor, G.K., 1991, Avalonian and Cadomian terranes: Precambrian Research, v. 53, p. 41–78, doi: 10.1016/0301-9268(91)90005-U. Nance, R.D., Murphy, J.B., and Keppie, J.D., 2002, A Cordilleran model for the evolution of Avalonia: Tectonophysics, v. 352, p. 11–31, doi: 10.1016/ S0040-1951(02)00187-7. Nardi, L.V.S., and Frantz, J.C., 1995, The Cordilheira Intrusive Suite: Late Proterozoic peraluminous granitoids from southern Brazil: Journal of South American Earth Sciences, v. 8, p. 55–63, doi: 10.1016/08959811(94)00041-Y. Nokleberg, W.J., Parfenov, L.M., Monger, J.W.H., Norton, I.O., Khanchuk, A.I., Stone, D.B., Scotese, C.R., Scholl, D.W., and Fujita, K., 2000, Phanerozoic tectonic evolution of the circum-North Pacific: Reston, Virginia, U.S. Geological Survey Professional Paper 1626, 122 p. O’Brien, S.J., O’Brien, B.H., Dunning, G.R., and Tucker, R.D., 1996, Late Neoproterozoic Avalonian and related peri-Gondwanan rocks of the Newfoundland Appalachians, in Nance, R.D., and Thompson, M.D., eds., Avalonian and related peri-Gondwanan terranes of the circum-North Atlantic: Boulder, Colorado, Geological Society of America Special Paper 304, p. 9–28. Paim, P.S.G., Chemale, F., Jr., and Lopes, R.C., 2000, A Bacia do Camaquã, in Holz, M., and De Ros, L.F., eds., Geologia do Rio Grande do Sul: Porto Alegre, Brazil, CIGO/Universidade Federal do Rio Grande do Sul, p. 251–274.
25
Patchett, P.J., Gale, N.H., Goodwin, R., and Humm, M.J., 1980, Rb-Sr whole rock isochron ages of late Precambrian to Cambrian igneous rocks from southern Britain: Journal of the Geological Society of London, v. 137, p. 649–656. Pedrosa-Soares, A.C., Noce, C.M., Wiedemann, C.M., and Pinto, C.P., 2001, The Araçuaí West-Congo orogen in Brazil: An overview of a confined orogen formed during Gondwanaland assembly: Precambrian Research, v. 110, p. 307–323, doi: 10.1016/S0301-9268(01)00174-7. Philipp, R.P., and Machado, R., 2001, Suítes graníticas do batólito Pelotas no Rio Grande do Sul: Petrografia, tectônica e aspectos petrogenéticos: Revista Brasileira de Geociências, v. 31, p. 257–266. Philipp, R.P., and Machado, R., 2005, The Late Neoproterozoic granitoid magmatism of the Pelotas Batholith, southern Brazil: Journal of South American Earth Sciences, v. 19, p. 461–478, doi: 10.1016/j.jsames.2005.06.010. Philipp, R.P., Mesquita, M.J., Gomes, M.E.B., and Almeida, D.P.M., 1993, Reconhecimento estrutural e geoquímico dos Granitóides Brasilianos da região de Pelotas: RS: Pesquisas, v. 20, p. 3–13. Pimentel, M.M., and Fuck, R.A., 1992, Neoproterozoic crustal accretion in central Brazil: Geology, v. 20, p. 375–379, doi: 10.1130/00917613(1992)020<0375:NCAICB>2.3.CO;2. Pimentel, M.M., Heaman, L., and Fuck, R.A., 1991, Zircon and sphene U-Pb geochronology of Upper Proterozoic volcanic-arc rock units from southwestern Goiás, central Brazil: Journal of South American Earth Sciences, v. 4, p. 295–305, doi: 10.1016/0895-9811(91)90003-4. Pimentel, M.M., Whitehouse, M.J., Viana, M.G., Fuck, R.A., and Machado, N., 1997, The Mara Rosa arc in the Tocantins Province: Further evidence for Neoproterozoic crustal accretion in central Brazil: Precambrian Research, v. 81, p. 299–310, doi: 10.1016/S0301-9268(96)00039-3. Pimentel, M.M., Fuck, R.A., and Botelho, N.F., 1999, Granites and the geodynamic history of the Neoproterozoic Brasília belt, central Brazil: A review: Lithos, v. 46, p. 463–483, doi: 10.1016/S0024-4937(98)00078-4. Porcher, C.C., and Fernandes, L.A.D., 1990, Relações embasamento-cobertura na porção occidental do Cinturão Dom Feliciano: Um esboço estrutural: Pesquisas, v. 17, p. 72–84. Preciozzi, F., Masquelin, H., and Basei, M.A.S., 1999, The Namaqua/Grenville terrane of eastern Uruguay, in II. South American Symposium on Isotope Geology, Cordoba, Argentina: Actas, p. 338–340. Ramos, V.A., 1988, Late Proterozoic–early Paleozoic of South America: A collisional history: Episodes, v. 11, p. 168–174. Remus, M.V.D., 1990. Geologia e geoquímica do Complexo Cambaizinho, São Gabriel, RS [M.S. thesis]: Porto Alegre, Brazil, Instituto de Geociências, Universidade Federal do Rio Grande do Sul, 267 p. Remus, M.V.D., Tedesco, M.A., Philipp, R.P., and Faccini, U.F., 1987, Evolução estrutural da unidade Porongos a sul do Rio Camaquã, RS, in Simpósio Sul-Brasileiro de Geologia 3rd, Curitiba: Anais Simpósio Sul-Brasileiro de Geologia 3: Sociedade Brasileira Geologia, v. 1, p. 223–244. Remus, M.V.D., Philipp, R.P., Faccini, U.F., and Junges, S.L., 1990, Contribução ao estudo geológico-estrutural dos Gnaisses Encantadas e as relações com as supracrustais Porongos na região de Santana da Boa Vista/RS, in Congresso Brasileiro de Geologia, 36th, Natal: Anais de Congresso Brasileiro de Geologia 36: SBG v. 5, p. 2356–2370. Remus, M.V.D., Hartmann, L.A., and Ribeiro, M., 1991, Nota sobre a geologia dos metamorfitos de pressão intermediária e granitóides associados da região de Pinheiro Machado/RS: Acta Geologica Leopoldensia, v. 34, p. 175–190. Remus, M.V.D., Hartmann, L.A., and Formoso, M.L.L., 1993, Os padrões de Elementos Terras Raras (ETR) e a afinidade geoquímica komatiítica dos xistos magnesianos e rochas associadas do Complexo Cambaizínho, São Gabriel-RS: Revista Brasileira de Geosciências, v. 23, no. 4, p. 370–387. Remus, M.V.D., McNaughton, N.J., Hartmann, L.A., Koppe, J.C., Fletcher, I.R., Groves, D.I., and Pinto, V.M., 1999, Gold in the Neoproterozoic juvenile Bossoroca volcanic arc of southernmost Brazil: Isotopic constraints to timing and sources: Journal of South American Earth Sciences, v. 12, p. 349–366, doi: 10.1016/S0895-9811(99)00026-7. Remus, M.V.D., Hartmann, L.A., McNaughton, N.J., Groves, D.I., and Fletcher, I.R., 2000a, The link between hydrothermal epigenetic copper mineralization and the Caçapava Granite of Brasiliano Cycle in southern Brazil: Journal of South American Earth Sciences, v. 13, p. 191–216, doi: 10.1016/S0895-9811(00)00017-1. Remus, M.V.D., Hartmann, L.A., McNaughton, N.J., Groves, D.I., and Reischl, J.L., 2000b, Distal magmatic-hydrothermal origin for the Camaquã Cu (Au-Ag) and Santa Maria Pb, Zn (Cu-Ag) deposits, southern Brazil: Gondwana Research, v. 3, p. 155–174, doi: 10.1016/S1342937X(05)70094-0.
26
Saalmann et al.
Remus, M.V.D., Macambira, M.B., Hartmann, L.A., and Beilfuss, M., 2001, Crustal evolution of granitoids and gneisses from the Cambaizinho belt, southern Brazil: Review, new zircon Pb-Pb evaporation ages and Pb-NdSr isotopes: III. South American Symposium on Isotope Geology, Pucon, Chile. CD-ROM. Rozendaal, A., Gresse, P.G., Scheepers, R., and Le Roux, J.N.P., 1999, Neoproterozoic to Early Cambrian crustal evolution of the Pan-African Saldania belt, South Africa: Precambrian Research, v. 97, p. 303–323, doi: 10.1016/S0301-9268(99)00036-4. Saalmann, K., Hartmann, L.A., Remus, M.V.D., Koester, E., and Conçeição, R.V., 2005a, Sm-Nd isotope geochemistry of metamorphic volcano-sedimentary successions in the São Gabriel block, southernmost Brazil: Evidence for the existence of juvenile Neoproterozoic oceanic crust to the east of the Rio de la Plata craton: Precambrian Research, v. 136, p. 159– 175, doi: 10.1016/j.precamres.2004.10.006. Saalmann, K., Remus, M.V.D., and Hartmann, L.A., 2005b, Geochemistry and crustal evolution of volcano-sedimentary successions and orthogneisses in the Sao Gabriel block, southernmost Brazil—Relics of Neoproterozoic magmatic arcs: Gondwana Research, v. 8, no. 2, p. 143–161, doi: 10.1016/S1342-937X(05)71114-X. Saalmann, K., Hartmann, L.A., and Remus, M.V.D., 2005c, Tectonic evolution of two contrasting schist belts in southernmost Brazil: A plate tectonic model for the Brasiliano orogeny: International Geology Review, v. 47, p. 1234–1259. Saalmann, K., Remus, M.V.D., and Hartmann, L.A., 2006, Structural evolution and tectonic setting of the Porongos belt, southern Brazil: Geological Magazine, v. 143, no. 1, p. 59–88, doi: 10.1017/S0016756805001433. Samson, S.D., Hibbard, J.P., and Wortman, G.L., 1995, Nd isotope evidence for juvenile crust in the Carolina terrane, southern Appalachians: Contributions to Mineralogy and Petrology, v. 121, p. 171–184, doi: 10.1007/ s004100050097. Samson, S.D., D’Lemos, R.S., Blichert-Toft, J., and Vervoort, J., 2003, U-Pb geochronology and Hf-Nd isotope compositions of the oldest Neoproterozoic crust within the Cadomian orogen: New evidence for a unique juvenile terrane: Earth and Planetary Science Letters, v. 208, p. 165–180, doi: 10.1016/S0012-821X(03)00045-1. Santos, J.O.S., Hartmann, L.A., Bossi, J., Campal, N., Schipilov, A., Piñeyro, D., and McNaughton, N.J., 2003, Duration of the Trans-Amazonian Cycle and its correlation within South America based on U-Pb SHRIMP geochronology of La Plata craton, Uruguay: International Geology Review, v. 45, p. 27–48. Scheepers, R., 1995, Geology, geochemistry and petrogenesis of Late Precambrian S-, I- and A-type granitoids in the Saldania belt, Western Cape Province, South Africa: Journal of African Earth Sciences, v. 21, p. 35–58, doi: 10.1016/0899-5362(95)00087-A. Schmitt, R.S., Trouw, R.A.J., Van Schmus, W.R., and Pimentel, M.M., 2004, Late amalgamation in the central part of West Gondwana: New geochronological data and the characterization of a Cambrian collisional orogeny in the Ribeira belt (SE Brazil): Precambrian Research, v. 133, p. 29–61, doi: 10.1016/j.precamres.2004.03.010. Seth, B., Kröner, A., Mezger, K., Nemchin, A.A., Pidgeon, R.T., and Okrusch, M., 1998, Archaean to Neoproterozoic magmatic events in the Kaoko belt of NW Namibia and their geodynamic significance: Precambrian Research, v. 92, p. 341–363, doi: 10.1016/S0301-9268(98)00086-2. Silva-Filho, B.C., and Soliani, E., Jr, 1987, Origem e evolucão dos Gnaisses Cambaí: Exemplo de estrudo integrado de análise estrutural, petroquímica e geocronologia. Simpósio Sul-Brasileiro de Geologia, 3rd, Curitiba, Brasil: Atas, v. 1, p. 127–145. Soliani, E., Jr., 1986, Os dados geocronológicos do Escudo Sul-rio-grandense e suas implicações de ordem geotectônica [Ph.D. thesis]: São Paulo, University of São Paulo, 425 p. Söllner, F., Lammerer, B., and Weber-Diffenback, K., 1989, Brasiliano age of a charnoenderbitic rock suite in the Complexo Costeiro (Ribeira Mobile Belt), Espírito Santo/Brazil: Evidence from U-Pb geochronology on zircons: Zentralblatt für Geologie und Paläontologie, Teil I, v. 5/6, p. 933–945. Söllner, F., Lammerer, B., and Weber-Diffenback, K., 1991, Die Krustenentwicklung nördlich von Rio de Janeiro/Brasilien: München, Münchner geologische Hefte, v. 4, 100 p. Stern, R.J., 1994, Arc assembly and continental collision in the Neoproterozoic east African orogen: Implications for the consolidation of Gondwanaland: Annual Review of Earth and Planetary Sciences, v. 22, p. 319–351.
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
<630 Ma (2) late orogenic plutons
Molassic deposits
Andesitic to rhyolitic lavas of intracontinental volcanic chain (580–550 Ma)
V1
V2
V3
V4
V5
Volcanic events
Lower P-II
Upper P-II
P-III
531 (5)
525 (6)
521–522 (7)
517 (8)
Age (Ma)
Regional nomenclature
Anti-Atlas Supergroup
Supergroup
Ouarzazate
Groups
Taroudant
and
Tata
Notes: Dating references: (1) Samson et al. (2004); (2) Thomas et al. (2004); (3) Gradstein et al. (2004); (4) Amthor et al. (2003); (5) Gasquet et al. (2005); (6) Maloof et al. (2005); (7) Compston et al. (1992), Landing et al. (1998); (8) Landing et al. (1998); (9) Tucker and McKerrow (1995).
630 (3)
Post-orogenic stage
Subvolcanic plutons
Local conglomerate deposits
Basaltic pile
Silts and shales
Dolostone
Thick limestone beds Silts, shales, volcanic conglomerate
S2
Interbedded limestones and shales
Thin limestone beds
Shales and calcarenites
Conglomerate
Fine sandstones and shales
C2
Precambrian-Cambrian boundary S1 (4)
Atdabanian
525
S3
C3
S4
Botomian
Cg
513
518
Cambrian
522
S5
Age (Ma) (3, 9)
Agoundis-Ounein formations
Adoudounian formations
TABLE 1. CHRONOSTRATIGRAPHY OF THE ATLASIC NEOPROTEROZOIC AND CAMBRIAN FORMATIONS
International subdivisions
Middle
Early Cambrian
Ediacaran
Cryogenian
30 Pouclet et al.
Geodynamic evolution of the northwestern Paleo-Gondwanan margin the age assignment and structural interpretation of the thick mafic volcanic pile overlying the Neoproterozoic rhyolitic complex and conformably overlain by the Cambrian deposits (the P-III basaltic volcanic rocks of Table 1), we perform a detailed study of this formation in two areas of the south and central Western HighAtlas, Agoundis-Ounein and Toubkal. It appears that this mafic pile is discordant above the rhyolitic complex and must be attributed to a distinct Late Ediacaran and Early Cambrian magmatic activity. A complete description of these volcanic rocks has been completed in the Agoundis-Ounein area. The structural features are fully determined in the Toubkal Massif. All these new data are used to propose a geodynamical evolution model for the Precambrian-Cambrian transition in the Moroccan Atlas. THE PRECAMBRIAN-CAMBRIAN BOUNDARY Structural Aspects A total of seven sites were investigated, at the edges of the Precambrian inliers (Fig. 1), where the Late Ediacaran and Early Cambrian transgressive formations overlay various older terranes. Lithological features are sketched in Figure 2. The regional nomenclature is displayed in Table 1. Site 1 is located east of the Kerdous inlier, in the western Anti-Atlas. The Precambrian formations are subhorizontal and belong to the Ediacaran formations of the Ouarzazate Supergroup, locally known as the Tanalt “Series” (Hassenforder, 1987). They consist of heterometric conglomerates made of subrounded centimeter- to decimeter-sized pebbles and blocks of quartz, rhyolites, and granites in an arkosic matrix. Locally, intercalations of acidic epiclastites are related to the underlying ignimbrite flows. The top of this formation passes gradually to a well-sorted conglomerate made of quartz, rhyolite, and granite pebbles that is attributed to the Adoudounian Basal Formation (Hassenforder, 1987). This formation, ~100 m thick, evolves upward to finer detrital deposits, with sandstones, silt, shales, and interbedded thin strata of siliceous limestones. It is overlain by the thick beds of the Tamjout dolostone. Almost the same succession is described around the other western Anti-Atlas inliers. However, to the east, the Basal Formation is decreasing in thickness and includes mafic volcanic products (Leblanc, 1977; Algouti et al., 2000, 2001; Benziane et al., 2002). Site 2 corresponds to the upper valley of the Agoundis River, west of Toubkal Mountain (the highest Atlas summit, culminating at 4167 m). The Precambrian formations are made of rhyolitic flows and pyroclastites, felsic epiclastites, sandstones, and conglomerates with centimeter-sized subrounded elements of rhyolites and granites, all overlain by a thick pile (more than 400 m) of basaltic flows. All these volcanic units are attributed to the local P-III Group (Proust, 1973), the upper part of the Ouarzazate Supergroup. The basaltic pile is conformably overlain by coarse-grained to conglomeratic sandstones, including basaltic pyroclastic fragments, quartz pebbles, and rhyolitic angular elements. The detrital sedimentation rapidly evolves upward to
31
finer sandstones, silts, and shales, with intercalated thin beds of siliceous dolostone. The latter sequences correspond to the Adoudounian Basal Formation defined throughout the western Anti-Atlas. Compared to the average feature of the Basal Formation (Choubert, 1963), the Agoundis equivalent is slightly thinner (140 m) and richer in fine-grained sediments (shales). It is overlain by the Tamjout dolostone. All the basaltic and overlying Adoudounian formations are gently dipping 30° to the west. Site 3 has been observed in central Anti-Atlas, south of Ouarzazate, close to the village of Anzel. The Precambrian formations consist of rhyolitic flows and pyroclastic deposits that dip 30° to the west. They belong to the local P-III of the Ouarzazate Supergroup. They are overlain, with a clear angular unconformity, by the Adoudounian Basal Formation, which is gently dipping 10° to the north. This formation, ~30 m thick, is made in its lower part of a conglomerate with well-rounded quartz pebbles in a pelitic matrix, and in its middle part of an alternation of microconglomerates, silty sandstones, and shales. The upper part is characterized by the occurrence of centimeter-sized interbedded siliceous dolostone layers, which become thicker and more frequent upward, until the dolostone pile of the Tamjout Formation. Sites 4 and 5 are located, respectively, north and southeast of the Bou Azzer inlier. At Site 4, the Precambrian deposits consist of a thick heterometric conglomerate with large pebbles of quartz, rhyolites, and granites, which is fairly dipping 15° to the east. They are attributed to the upper part of the Ouarzazate Supergroup (P-III). The slightly unconformably upper-lying Adoudounian terrain is subhorizontal and shows the same lithology as that of Site 3. At Site 5, the Precambrian formations are quite similar, but are more steeply dipping 50° to the NNE and 40° to the SSW, on each side of a WNW–ESE-trending anticlinal fold axis. The conglomerate includes large blocks of granites and pebbles of rhyolites, ignimbrites, and andesites. It overlies a succession of sandstones and siltstones in a typical reverse-graded bedding sequence. All these deposits are gathered in the local Trifya Formation that we attributed to the Ouarzazate Supergroup, and not to the Anti-Atlas Supergroup, as is commonly done by attribution of the Trifya sediments to the Tiddiline “Series” (Hefferan et al., 1992). This difference in the chronostratigraphic attribution is important, because the Tiddiline formations registered a penetrative deformation having the structural characteristics of the major Pan-African tectonic event, and thus they are included in the Anti-Atlas Supergroup dated to the Cryogenian Period (Table 1). If the Trifya formations are correlated with the Tiddiline ones, the Cambrian unconformity is explained by the Pan-African event, which predated the Ouarzazate Supergroup setting. But in the Bou Azzer inlier, the Tiddiline and Trifya sedimentary series are located in two tectonically separated basins. The Trifya formations are not affected by the Pan-African penetrative deformation. Besides, the lithological features of the Tiddiline “Series” (diamictite conglomerates, thick green claystones including ice-rafting quartz pebbles, and interbedded mafic lavas) are totally lacking in the Trifya formations that are constituted only by fine- to coarse-grained siliciclastic sediments. Moreover,
Cambrian
SE-Bou Azzer
N-Bou Azzer
conglomerate
6
Rhyolite flows and pyroclastites
Basal Formation (30 m)
Tamjout dolostone
Rhyolite-andesite complex
Detrital complex (200 m)
Middle Cambrian
Figure 2. Lithological logs of the seven selected Precambrian-Cambrian boundary sites.
Rhyolite-andesite complex
Trifya Formation
Basal Formation (30 m)
Tamjout dolostone
5
Rhyolitic complex
Basalt flows (>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
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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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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
50
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 <3
28 2145.0 75.5 1977.0 0.94 5 0.51 1 27
32 2698.1 58.6 2465.3 1.41 11 <0.05 <0.5 95
27 1822.4 90.9 1646.0 0.64 59 0.07 <0.5 15
43 3783.0 106.2 1567.4 1.75 4 0.52 <0.5 13
93 2363.0 100.0 833.0 3.27 9 2.61 1 11
149 779.4 40.3 126.0 6.41 59 3.74 3 56
La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
0.169 0.481 0.091 0.450 0.159 0.050 0.204 0.040 0.230 0.057 0.183 0.026 0.172 0.029
0.926 1.960 0.207 0.811 0.144 0.050 0.115 0.020 0.103 0.025 0.065 0.008 0.054 0.008
0.296 0.413 <0.040 0.237 <0.060 <0.020 <0.070 <0.010 <0.050 <0.010 <0.040 <0.010 <0.030 <0.010
0.228 0.395 <0.040 <0.150 <0.060 <0.020 <0.070 <0.010 <0.050 <0.010 <0.040 <0.010 <0.030 <0.010
<0.050 <0.050 <0.040 <0.150 <0.060 <0.020 0.078 <0.010 0.071 <0.010 0.062 <0.010 0.071 <0.010
0.108 0.313 0.065 0.377 0.233 0.097 0.315 0.058 0.369 0.093 0.243 0.041 0.276 0.043
0.133 0.481 0.090 0.483 0.270 0.108 0.426 0.084 0.602 0.140 0.386 0.063 0.461 0.059
(wt%)
Note: See text. Analytical laboratory of the Research Center for Petrography and Geochemistry of Nancy, France. Same method as for Table 2. LOI—loss on ignition.
observed. A submeridian compressional tectonic event in the latest Neoproterozoic has to be considered. We analyzed the volcanic activities at the Precambrian-Cambrian transition time, and we emphasize the significantly different magmatic signatures between the Ediacaran and the Late Ediacaran–Early Cambrian volcanic rocks. The Ediacaran lavas are andesites to rhyolites of a volcanic chain emplaced in an intracontinental setting. They were generated by melting of a water- and lithophile element–enriched mantle during previous subduction events in the eastern to central Anti-Atlas and in the south of High-Atlas, and by crustal assimilation and crustal melting throughout the Atlas
area. A present-time equivalent could be Anatolia in Turkey. The Late Ediacaran and Early Cambrian lavas are continental tholeiites that outpoured under shallow water at a marine basin margin. South of the western High-Atlas, they constitute a basaltic pile, associated with a dike swarm trending N30°. They were generated by melting of the transitional lithospheric mantle. The report of the different magmatic signatures gains a particular geotectonic significance with the discovery of a major structural unconformity between the upper rhyolitic formations of the Ediacaran Atlasic Volcanic Chain and the basaltic pile allotted to the Adoudounian Basal Formation. Indeed, the basalts
Geodynamic evolution of the northwestern Paleo-Gondwanan margin Ounein lherzolite Sidi Flah lherzolite Bou Azzer pyroxenite Bou Azzer wherlite
Rock / Chondrite CI
10.00
55
Galicia Margin peridotites
Figure 13. Rare earth element patterns normalized to chondrite for the ultramafic rocks of Ounein, Sidi Flah, and Bou Azzer. Galicia Margin, compositional domain of selected peridotites having high MgO and REE contents, after Seifert and Brunotte (1996). Normalization values after Evensen et al. (1978). The Ounein lherzolite is less depleted than the Galicia Margin peridotites.
1.00
0.10
La
Ce
Pr
Nd
Pm
Sm
Eu
Gd
Tb
Dy
Ho
Er
were formerly considered as conformably overlying the andesitic and rhyolitic lavas and produced in a banal magmatic evolution from calc-alkaline to tholeiitic. A tectonomagmatic continuity from Precambrian to Cambrian was inferred. This postulate has to be revised. Precambrian to Cambrian Geodynamic Evolution of the South Atlasic Region Taking into account the new data presented above, we propose the following geodynamic evolution at the PrecambrianCambrian boundary in the South Atlasic region, located at the northwestern border of Paleo-Gondwana. This Paleo-Gondwana margin was constituted by amalgamation of eastern Anti-Atlas, High-Atlas, and Meseta terranes toward the west African craton as a consequence of the Pan-African orogeny. In the Anti-Atlas, the late orogenic period ended with deposition of coarse molassic conglomerates. Then the Ediacaran Atlasic Volcanic Chain, an intracontinental andesitic to rhyolitic volcanic chain, covered the entire Anti-Atlas and part of the High-Atlas. The tectonic context is transtensional, and a mantle upwelling can be suspected (Gasquet et al., 2005). This tectonomagmatic event has to be placed in the geodynamic evolution of the Avalonian realm. At that time (570–550 Ma), an intracontinental wrench regime developed in the Avalonian region, which could be linked to ridge-trench collision and the formation of a major transform fault system, according to Murphy et al. (2000) and Nance et al. (2002). Cessation of the main phase of the Avalonian magmatism is explained by the termination of subduction. Indeed, the building of the Ediacaran Atlasic Volcanic Chain coincided with the cessation of the Avalonian activity and, thus, with the suspected ridge convergence. In their “Baja” model, Keppie et al. (2003) show that the “Merlin-Morgana” ridge went underneath
Tm
Yb
Lu
the Atlasic region between 570 and 550 Ma (Fig. 14). This event would be an opportunity for supplying heat, which was channeled along the Pan-African inherited fractures. This heat transfer is necessary to melt the contaminated upper mantle and the lower crust that produced the Ediacaran Atlasic Volcanic Chain magmas. The subduction regime of the Paleo-Gondwana margin evolved to a transform regime, which caused a right-lateral motion along the west African craton (Fernandez-Suarez et al., 2002; Gutiérrez-Alonso et al., 2003), particularly along the south Atlasic tectonic zone. Hence, the necessary structural and thermal conditions are present to produce the volcanic chain. For many authors, the Late Neoproterozoic tensional regime continued with the Early Cambrian extension and marine transgression. However, as shown above, a compressional event occurred in the latest Neoproterozoic, which caused folding of the Ouarzazate Supergroup, north of the Anti-Atlas major fault, and probable lateral displacements along the Tizi n’Test and South Atlas fault system (TNT and SAF, Fig. 1). For structural (unconformity between the Ediacaran Atlasic Volcanic Chain and the overlying basalts) and magmatologic (different magmatic sources) reasons, the Late Ediacaran–Early Cambrian tholeiites (V1 and V2) must be attributed to a new geotectonic setting. They are linked to a WNW-ESE true extension, pointed out by northwest-facing northeast-southwest normal faults and by a N30° fissural system (Fig. 15). This tectonic feature, the associated transgression from a new marine basin, the continental tholeiite magmatic signature, and the presence of lherzolite fragments of transitional mantle origin strongly suggest a volcanic passive margin context. Although set farther to the east, the Djbel Boho magmatic activity, showing an OIB-type mantle source, indicates a subcontemporaneous asthenosphere upwelling. However, if the continental margin was thinned, with local mantle denudation, remnant continental crust underlies the marine basin, as shown
56
Pouclet et al.
rid
ge Cd Ib
Ms Cr
EAVC
WA
WAC Sg Fl Yu Figure 14. Late Neoproterozoic ridge-trench collision to subduction at the border of the Avalonian terranes, in the “Baja” model of Keppie et al. (2003). To the known Avalonian-Cadomian terranes—Cadomia (Cd), Carolina (Cr), Florida (Fl), Iberia (Ib), West Avalonia (WA), and Yucatan (Yu)—one must add the Meseta (Ms) and Sénégal blocks (Sg), which remained attached to the African plate after the early Paleozoic drifting. EAVC—Ediacaran Atlasic Volcanic Chain; WAC—West African craton.
8°
schem
32°
atic cr
oss-se
edg rift
TNT
Eas tern
Wes tern grab en
es
ction
F
by the lower crust xenoliths collected in the late Early and Middle Cambrian volcanic rocks. The extension was limited, because no more lava with asthenospheric signature appeared in the basin edges. The rifting rapidly aborted and no mid-ocean ridge basalt (MORB)-type lavas were set in the Moroccan Cambrian basin. The deepest graben is located on the western side of the Meseta, where more than 7000 m of Cambrian terrigenous sediments are deposited (Fig. 15). It is controlled by N30° and N60° faults (Bernardin et al., 1988). The overall trend is SSW-NNE, but because of the Variscan strong reverse folding and eastward thrust that trends N20°, the original shape of the trough edges is poorly determined. The true ocean opening probably took place farther to the northwest, between Meseta and the rest of Avalonian blocks. Contrary to the assumption of Landing (1996), there are no drastic facies differences between Avalon and Morocco, which cannot be explained by local sedimentary conditions. The lithostratigraphic column cited by Landing (1996) concerns the eastern shallow water part of the Cambrian basin, and not the deep water ones, which are exposed at the westernmost side of HighAtlas (de Koning, 1957). Besides, the faunal differences from the Atdabanian (Geyer, 1998; comments of Landing, 1996) are consistent with a rifting of Avalon from Paleo-Gondwana. Similarly, in the Avalon continental margin, one can observe small Early Cambrian grabens containing the same continental tholeiites as in the Atlasic border (Greenough et al., 1985; Murphy et al., 1985). The last question is: what happened ca. 545 Ma that was responsible for cessation of the activity in the Ediacaran Atlasic Volcanic Chain, local compression, and intracontinental rifting, leading to drifting of the previously accreted Avalonian terranes? One possibility is that the subduction of Paleo-Pacific ridges drastically modified the disposition of asthenosphere convection cells and placed the uprising mantle below the former suture zones. The Early Cambrian rifting event has to be explained by reorganization of the plate-scale motions, in relationship with the vanishing of the Avalonian-sided oceans and the extension-subduction process of the Paleo-Pacific domain.
SA
CONCLUSIONS
dy k sw e arm
AAMF
Db 30° 8°
Figure 15. Early Cambrian rifting in the west Atlas region. Dike swarm of the Toubkal Massif. Abbreviations are as in Figure 1. The cross-section of the rift shoulders is reconstituted at the time before the Variscan eastward thrust and the Atlasic tectonic uplift of the High-Atlas. The basaltic activity of the Agoundis-Ounein and Toubkal area is located at the eastern edges of the rift; it took place in the Meseta and to the west of the High-Atlas and the Anti-Atlas (Bernardin et al., 1988; Piqué et al., 1995).
At the Late Neoproterozoic–Cambrian boundary, a succession of tectonic and magmatic events occurred at the northwestern margin of the Paleo-Gondwana in the Moroccan Atlas domain. Particular attention is given to the tectonic conditions that prevailed during the Cambrian marine transgression and to the tectonomagmatic significance of the volcanic activities, just before and just after the Late Ediacaran–Early Cambrian geodynamic climax. Major structural unconformities are demonstrated between the Ediacaran calc-alkaline volcanic complex and the Late Ediacaran and Early Cambrian flood basalts. An intracontinental volcanic chain (Ediacaran Atlasic Volcanic Chain) was built between 580 and 550 Ma. This N65°-trending chain reaches a length of 850 km and straddles the Pan-African orogenic area.
Geodynamic evolution of the northwestern Paleo-Gondwanan margin It covers the western substratum of the west African craton, the central Anti-Atlas suture zone, and the eastern Anti-Atlas mobile belt. The main Pan-African orogenic stage occurred ca. 660 Ma (Thomas et al., 2002), together with amalgamation of the Avalonian terranes (Nance et al., 2002). Subsequently the Ediacaran Atlasic Volcanic Chain activity took place in a postorogenic transtensional context, reworking the Pan-African fracture zones. The magma genesis, with melting of contaminated mantle of the previous mobile belt and melting of the crust, is consistent with this setting. Cessation of this activity was followed by a local NNE-SSW compression and then by a WNW-ESE extension. In the old block of High-Atlas, central Anti-Atlas, and eastern AntiAtlas, the Late Ediacaran and Cambrian transgressive formations unconformably overlie the Ediacaran Atlasic Volcanic Chain volcanic rocks. In the High-Atlas and central Anti-Atlas areas, the sequence begins with outpouring of flood basalts of continental tholeiite composition, which were generated by melting of the subcontinental lithospheric mantle. Structural, magmatic, and petrographic (lherzolite material) data favor a volcanic passive margin geotectonic setting. However, the Early Cambrian rifted basin did not evolve to an ocean basin. Mantle denudation was not complete, magmatic activity changes from tholeiitic to alkaline, with a lesser degree of partial melting, and there were no depleted MORB-type basalts. ACKNOWLEDGMENTS This study was supported by the Institut des Sciences de la Terre of the University of Orléans, France, and by the FrenchMoroccan Action Intégrée 222/STU/00. We thank Jean-Paul Liégeois and Nasser Ennih for critical reviews of the first version of the manuscript. REFERENCES CITED Admou, H., 2000, Structuration de la paléosuture ophiolitique panafricaine de Bou Azzer-Siroua (Anti-Atlas central, Maroc) [Ph.D. thesis]: Marrakech, Morocco, University of Cadi Ayyad, 201 p. Admou, H., and Juteau, T., 1998, Découverte d’un système hydrothermal océanique fossile dans l’ophiolite antécambrienne de Khzama (massif du Siroua, Anti-Atlas marocain): Comptes Rendus Académie des Sciences, Paris, Sciences de la terre et des planètes, v. 327, p. 335–340. Aït Malek, H., Gasquet, D., Bertrand, J.M., and Leterrier, J., 1998, Géochronologie U-Pb sur zircon de granitoïdes éburnéens et panafricains dans les boutonnières protérozoïques d’Igherm, du Kerdous et du Bas Drâa (AntiAtlas occidental, Maroc): Comptes Rendus Académie des Sciences, Paris, Sciences de la terre et des planètes, v. 327, p. 819–826. Algouti, Ab., Algouti Ah., Beauchamp, J., Chbani, B., and Taj-Eddine, K., 2000, Paléogéographie d’une plateforme infracambrienne de la région Waoufengha-Igherm, Anti-Atlas, Maroc: Comptes Rendus Académie des Sciences, Paris, Sciences de la terre et des planètes, v. 330, p. 155–160. Algouti, Ab., Algouti, Ah., Chbani, B., and Zaim, M., 2001, Sédimentation et volcanisme synsédimentaire de la série de base de l’adoudounien infracambrien à travers deux exemples de l’Anti-Atlas du Maroc: Journal of African Earth Sciences, v. 32, p. 541–556. Alvaro, J.J., Ezzouhairi, H., Vennin, E., Ribeiro, M.L., Clausen, S., Charif, A., Ait Ayad, N., and Moreira, M.E., 2006, The Early-Cambrian Boho volcano of the El Graara massif, Morocco: Petrology, geodynamic setting and coeval sedimentation: Journal of African Earth Sciences, v. 44, p. 396–410.
57
Amthor, J., Grotzinger, J., Schröder, S., Bowring, S., Ramezani, J., Martin, M., and Matter, A., 2003, Extinction of Cloudina and Namacalathus at the Precambrian-Cambrian boundary in Oman: Geology, v. 31, p. 431–434, doi: 10.1130/0091-7613(2003)031<0431:EOCANA>2.0.CO;2. Azizi-Samir, M.R., Ferrandini, J., and Tane, J.-L., 1990, Tectonique et volcanisme tardi-Pan Africains (580–560 M.a.) dans l’Anti-Atlas Central (Maroc): Interprétation géodynamique à l’échelle du NW de l’Afrique: Journal of African Earth Sciences, v. 10, p. 549–563, doi: 10.1016/0899-5362(90)90105-N. Badra, L., Pouclet, A., Prost, A., and Touray, J.-C., 1992, Mise en évidence d’une extension intra-plaque tardi-panafricaine d’intérêt métallogénique dans le Haut Atlas occidental (Maroc): Comptes Rendus Académie des Sciences, Paris, v. 314, p. 703–709. Bajja, A., 2001, Volcanisme syn- à post-orogénique du Néoprotérozoïque de l’Anti-Atlas. Implications pétrogénétiques et géodynamiques, in De Wall, H., and Greiling, R.O., eds, Magmatic evolution of a Neoproterozoic island-arc: Syn- to post-orogenic igneous activity in the Anti-Atlas (Morocco): Jülich, Germany, Forschungszentrum Jülich, v. 45, p. 9–227. Benammi, M., Toto, El A., and Chakiri, S., 2001, Les chevauchements frontaux du Haut Atlas central marocain: Styles structuraux et taux de raccourcissement différentiel entre les versants nord et sud: Comptes Rendus Académie des Sciences, Paris, Sciences de la terre et des planètes, v. 333, p. 241–247. Benssaou, M., and Hamoumi, N., 2003, Le graben de l’Anti-Atlas occidental (Maroc): Contrôle tectonique de la paléogéographie et des séquences au Cambrien inférieur: Comptes Rendus Géoscience, v. 335, p. 297–305, doi: 10.1016/S1631-0713(03)00033-6. Benziane, F., Prost, A.E., and Yazidi, A., 1983, Le passage du Précambrien au Cambrien précoce volcanique et sédimentaire de l’Anti-Atlas oriental; comparaisons avec l’Anti-Atlas occidental: Bulletin Société Géologique de France, v. 7, t. XXV, p. 549–556. Benziane, F., Yazidi, A., Walsh, G.J., Armstrong, T.R., Kouhen, M.A., Yazidi, M., El Khamlichi, M.A., and Aleinikoff, J.N., 2002, Carte géologique du Maroc au 1/50 000. Feuille Zawyat Si Nisser. Mémoire explicatif: Notes et Mémoires, Service Géologique du Maroc, v. 423 bis, 69 p. Beraaouz, E.H., Ikenne, M., Mortaji, A., Madi, A., Lahmam, M., and Gasquet, D., 2004, Neoproterozoic granitoids associated with the Bou-Azzer ophiolitic melange (Anti-Atlas, Morocco): Evidence of adakitic magmatism in an arc segment at the NW edge of the west-African craton: Journal of African Earth Sciences, v. 39, p. 285–293, doi: 10.1016/j.jafrearsci.2004.07.040. Bernardin, Ch., Cornée, J.-J., Corsini, M., Mayol, S., Muller, J., and Tayebi, M., 1988, Variations d’épaisseur du Cambrien moyen en Meseta marocaine occidentale: Signification géodynamique des données de surface et de subsurface: Canadian Journal of Earth Sciences, v. 25, p. 2104–2117. Boudda, A., and Choubert, G., 1972, Sur la limite inférieure du Cambrien au Maroc: Comptes Rendus Académie des Sciences, Paris, v. 275, p. 5–8. Bouougri, E.H., 2003, The Moroccan Anti-Atlas: The West African craton passive margin with limited Pan-African activity. Implications for the northern limit of the craton: Precambrian Research, v. 120, p. 179–183, doi: 10.1016/S0301-9268(02)00169-9. Bouougri, E.H., and Saquaque, A., 2004, Lithostratigraphic framework and correlation of the Neoproterozoic northern West African craton passive margin sequence (Siroua-Zenaga-Bouazzer Elgraara Inliers, Central AntiAtlas, Morocco): An integrated approach: Journal of African Earth Sciences, v. 39, p. 227–238, doi: 10.1016/j.jafrearsci.2004.07.045. Charlot, R., 1982, Caractérisation des événements éburnéens et panafricains dans l’Anti-Atlas marocain: Apport de la méthode géochronologique RbSr: Notes et Mémoires, Service Géologique du Maroc, v. 313, p. 1–106. Chbani, B., Beauchamp, J., Algouti, A., and Zouhair, A., 1999, Un enregistrement sédimentaire éocambrien dans un bassin intracontinental en distension: Le cycle «conglomérat de base—unité calcaire—grès de Tikirt» de Bou-Azzer El Graara (Anti-Atlas central, Maroc): Comptes Rendus Académie des Sciences, Paris, Sciences de la terre et des planètes, v. 329, p. 317–323. Cheilletz, A., Levresse, G., Gasquet, D., Azizi-Samir, M.R., Zyadi, R., Archibald, D.A., and Farrar, E., 2002, The giant Imiter silver deposit: Neoproterozoic epithermal mineralization in the Anti-Atlas, Morocco: Mineralium Deposita, v. 37, p. 772–781, doi: 10.1007/s00126-002-0317-0. Choubert, G., 1963, Histoire géologique du Précambrien de l’Anti-Atlas: Notes et Mémoires, Service Géologique du Maroc, v. 162, 352 p. Compston, W., Williams, I.S., Kirschvink, J.L., Zichao, Z., and Guogan, M., 1992, Zircon U-Pb ages for the Early Cambrian time-scale: Journal of the Geological Society, London, v. 149, p. 171–184.
58
Pouclet et al.
Debrenne, F., and Debrenne, M., 1978, Archaeocyathid fauna of the lowest fossiliferous levels of Tiout (Lower Cambrian, Southern Morocco): Geological Magazine, v. 115, p. 101–119. de Koning, G., 1957, Géologie des Ida Ouzal (Maroc). Stratigraphie, pétrographie et tectonique de la partie SW du compartiment occidental du massif ancien du Haut Atlas: Leidse Geologische Mededelingen, Holland, v. 23, 209 p. Demange, M., 1980, Stratigraphie, volcanisme et paléogéographie du Précambrien III et de la série de base dans la partie sud de la boutonnière d‘Ouaoufengha (Anti-Atlas occidental) et minéralisation associée: Notes et Mémoires, Service Géologique du Maroc, v. 41, p. 7–23. Dick, H.J.B., and Bullen, T., 1984, Chromian spinel as a petrogenetic indicator in abyssal and alpine-type peridotites and spatially associated lavas: Contributions to Mineralogy and Petrology, v. 86, p. 54–76, doi: 10.1007/ BF00373711. Ducrot, J., and Lancelot, J.R., 1977, Problème de la limite Précambrien-Cambrien; étude radiochronologique par la méthode U-Pb sur zircons du volcan du Jbel Boho (Anti-Atlas marocain): Canadian Journal of Earth Sciences, v. 14, p. 2771–2777. Duke, J.M., 1983, Ore deposit model 7: Magmatic segregation deposits of chromite: Geosciences, Canada, v. 10, p. 15–24. El Archi, A., El Houicha, M., Jouhari, A., and Bouabdelli, M., 2004, Is the Cambrian basin of the Western High Atlas (Morocco) related either to a subduction zone or a major shear zone? Journal of African Earth Sciences, v. 39, p. 311–318, doi: 10.1016/j.jafrearsci.2004.07.037. El Attari, A., Hoepffner, C., and Jouhari, A., 1997, Nouvelles données magmatiques et structurales en relation avec la cinématique de l’ouverture du bassin cambrien de la Méseta occidentale (Maroc): Gaia, Lisboa, v. 14, p. 11–21. El Boukhari, A., Chabane, A., Rocci, G., and Tane, J.-L., 1992, Upper proterozoic ophiolites of the Siroua Massif (Anti-Atlas, Morocco): A marginal sea and transform fault system: Journal of African Earth Sciences, v. 14, p. 67–80, doi: 10.1016/0899-5362(92)90056-I. El Hadi, H., Tahiri, A., Simancas Cabrera, F., Gonzales Lodeiro, F., Azor Pérez, A., and Martinez Poyatos, D.J., 2006, Un exemple de volcanisme calcoalcalin de type orogénique mis en place en contexte de rifting (Cambrien de l’oued Rhebar, Meseta occidental, Maroc): Comptes Rendus Géoscience, v. 338, p. 229–236, doi: 10.1016/j.crte.2005.12.006. Ennih, N., and Liégeois, J.-P., 2001, The Moroccan Anti-Atlas: The West African craton passive margin with limited Pan-African activity. Implications for the northern limit of the craton: Precambrian Research, v. 112, p. 289– 302, doi: 10.1016/S0301-9268(01)00195-4. Ennih, N., and Liégeois, J.-P., 2003, The Moroccan Anti-Atlas: The West African craton passive margin with limited Pan-African activity. Implications for the northern limit of the craton: Reply to comments by E.H. Bouougri: Precambrian Research, v. 120, p. 185–189, doi: 10.1016/S03019268(02)00168-7. Evans, C.A., and Girardeau, J., 1988, Galicia Margin peridotites: Undepleted abyssal peridotites from the North Atlantic, in Boillot, G., Winterer, E.L., et al., eds, Proceedings of ODP, scientific results: College Station, Texas, Ocean Drilling Program, v. 103, p. 195–207. Evensen, N.M., Hamilton, P.J., and O’Nions, R.K., 1978, Rare earth abundances in chondritic meteorites: Geochimica et Cosmochimica Acta, v. 42, p. 1199–1212, doi: 10.1016/0016-7037(78)90114-X. Ezzouhairi, H., 2001, Le magmatisme post-collisionnel panafricain (tardi à post-orogénique) des régions d’Aghbalou, Sidi Flah-Bouskour et Oued Imini (Ouarzazate, Anti Atlas central, Maroc). Lithostratigraphie, géochimie, pétrogenèse et contexte géodynamique [Ph.D. thesis]: El Jadida, Morocco, University Chouaïb Doukhali, 201 p. Fekkak, A., Pouclet, A., and Benharref, M., 2003, The Middle Neoproterozoic Sidi Flah Group (Anti-Atlas, Morocco): Synrift deposition in a Pan-African continent-ocean transition zone: Journal of African Earth Sciences, v. 37, p. 73–87, doi: 10.1016/S0899-5362(03)00049-6. Fernandez-Suarez, J., Gutiérez Alonso, G., and Jeffries, T.E., 2002, The importance of along-margin terrane transport in northern Gondwana: Insights from detrital zircon parentage in Neoproterozoic rocks from Iberia and Brittany: Earth and Planetary Sciences Letters, v. 204, p. 75–88, doi: 10.1016/S0012-821X(02)00963-9. Gasquet, D., Levresse, G., Cheilletz, A., Azizi-Samir, M.R., and Mouttaqi, A., 2005, Contribution to a geodynamic reconstruction of the AntiAtlas (Morocco) during Pan-African times with the emphasis on inversion tectonics and metallogenic activity at the Precambrian-Cambrian
transition: Precambrian Research, v. 140, p. 157–182, doi: 10.1016/ j.precamres.2005.06.009. Geyer, G., 1998, Intercontinental, trilobite-based correlation of the Moroccan early Middle Cambrian: Canadian Journal of Earth Sciences, v. 35, p. 374–401, doi: 10.1139/cjes-35-4-374. Gradstein, F.M., Ogg, J.G., Smith, A.G., Bleeker, W., and Lourens, L.J., 2004, A new geologic time scale, with special reference to Precambrian and Neogene: Episodes, v. 27, p. 83–100. Greenough, J.D., McCurcheon, S.R., and Papezik, V.S., 1985, Petrology and geochemistry of Cambrian volcanic rocks from the Avalon Zone in New Brunswick: Canadian Journal of Earth Sciences, v. 22, p. 881–892. Gutiérrez-Alonso, G., Fernandez-Suarez, J., Jeffries, T.E., Jenner, G.A., Tubrett, M.N., Cox, R., and Jackson, S.E., 2003, Terrane accretion and dispersal in the northern Gondwana margin. An Early Paleozoic analogue of a long-lived active margin: Tectonophysics, v. 365, p. 221–232, doi: 10.1016/S0040-1951(03)00023-4. Hassenforder, B., 1987, La tectonique panafricaine et varisque de l’Anti-Atlas dans le massif de Kerdous (Maroc) [Ph.D. thesis]: France, University of Strasbourg, 249 p. Hefferan, K.P., Karson, J.A., and Saquaque, A., 1992, Proterozoic collisional basins in a Pan-African suture zone, Anti-Atlas Mountains, Morocco: Precambrian Research, v. 54, p. 295–319, doi: 10.1016/0301-9268(92)90075-Y. Hefferan, K.P., Admou, H., Karson, J.A., and Saquaque, A., 2000, Anti-Atlas (Morocco) role in Neoproterozoic western Gondwana reconstruction: Precambrian Research, v. 103, p. 89–96, doi: 10.1016/S03019268(00)00078-4. Hefferan, K.P., Admou, H., Hilal, R., Karson, J.A., Saquaque, A., Juteau, T., Bohn, M.M., Samson, S.D., and Kornprobst, J.M., 2002, Proterozoic blueschist-bearing melange in the Anti-Atlas Mountains, Morocco: Precambrian Research, v. 118, p. 179–194, doi: 10.1016/S0301-9268(02)00109-2. Hickey, R.L., Frey, F.A., and Gerlach, D.C., 1986, Multiple sources for basaltic arc rocks from the southern volcanic zone of the Andes (34–41°S): Trace element and isotopic evidence for contributions from subducted oceanic crust, mantle and continental crust: Journal of Geophysical Research, v. 91, p. 5963–5983. Hoepffner, C., Soulaimani, A., and Piqué, A., 2005, The Moroccan Hercynides: Journal of African Earth Sciences, v. 43, p. 144–165, doi: 10.1016/ j.jafrearsci.2005.09.002. Hoepffner, C., Houari, M.R., and Bouabdelli, M., 2006, Tectonics of the North African Variscides (Morocco, western Algeria): An outline: Comptes Rendus Géoscience, v. 338, p. 25–40, doi: 10.1016/j.crte.2005.11.003. Holm, P.E., 1985, The geochemical fingerprints of different tectono-magmatic environments using hygromagmatophile element abundance of tholeiitic basalts and basaltic andesites: Chemical Geology, v. 51, p. 303–323, doi: 10.1016/0009-2541(85)90139-1. Houari, M.R., and Hoepffner, C., 2003, Late Carboniferous dextral wrenchdominated transpression along the North African craton margin (Eastern High-Atlas, Morocco): Journal of African Earth Sciences, v. 37, p. 11–24, doi: 10.1016/S0899-5362(03)00085-X. Hupé, P., 1952, Contribution à l’étude du Cambrien inférieur et du Précambrien III de l’Anti-Atlas marocain: Notes et Mémoires, Service Géologique du Maroc, v. 103, p. 1–402. Inglis, J.D., MacLean, J.S., Samson, S.D., D’Lemos, R.S., Admou, H., and Hefferan, K., 2004, A precise U-Pb zircon age for the Bleïda granodiorite, Anti-Atlas, Morocco: Implications for the timing of deformation and terrane assembly in the eastern Anti-Atlas: Journal of African Earth Sciences, v. 39, p. 277–283, doi: 10.1016/j.jafrearsci.2004.07.041. Irvine, T.N., 1967, Chromian spinel as a petrogenetic indicator. Part 2, Petrologic applications: Canadian Journal of Earth Sciences, v. 2, p. 648–672. Jouhari, A., El-Archi, A., Aarab, M., El-Attari, A., Ennih, N., and Laduron, D., 2001, Géochimie et cadre géodynamique du volcanisme néoprotérozoïque terminal (vendien) du Haut Atlas occidental, Maroc: Journal of African Earth Sciences, v. 32, p. 695–705, doi: 10.1016/S0899-5362(02)00049-0. Keppie, J.D., Nance, R.D., Murphy, J.B., and Dostal, J., 2003, Tethyan, Mediterranean, and Pacific analogues for the Neoproterozoic-Paleozoic birth and development of peri-Gondwana terranes and their transfer to Laurentia and Laurussia: Tectonophysics, v. 365, p. 195–219, doi: 10.1016/ S0040-1951(03)00037-4. Kornprobst, J., Ohnenstetter, P., and Ohnenstetter, M., 1981, Na and Cr contents in clinopyroxene from peridotites: A possible discrimination between “subcontinental” and “suboceanic” mantle: Earth and Planetary Science Letters, v. 53, p. 241–254, doi: 10.1016/0012-821X(81)90157-6.
Geodynamic evolution of the northwestern Paleo-Gondwanan margin Kürkçüoglu, B., Sen, E., Aydar, E., Gourgaud, A., and Gündogdu, N., 1998, Geochemical approach to magmatic evolution of Mt Erciyes stratovolcano, Central Anatolia, Turkey: Journal of Volcanology and Geothermal Research, v. 85, p. 473–494, doi: 10.1016/S0377-0273(98)00067-5. Landing, E., 1996, Avalon: Insular continent by the past Precambrien, in Nance, R.D., and Thompson, M.D., eds, Avalonian and related peri-Gondwanan terranes of the circum-North Atlantic: Boulder, Colorado, Geological Society of America Special Paper 304, p. 29–63. Landing, E., Bowring, S.A., Davidek, K.L., Westrop, S.R., Geyer, G., and Heldmaier, W., 1998, Duration of the Early Cambrian: U-Pb ages from Avalon and Gondwana: Canadian Journal of Earth Sciences, v. 35, p. 329–338, doi: 10.1139/cjes-35-4-329. Leblanc, M., 1977, Synchronisme des faciès volcaniques (Précambrien III) et sédimentaires (Adoudounien) dans l’Infracambrien d’Alous (Anti-Atlas, Maroc): Comptes Rendus Académie des Sciences, Paris, v. 284, D, p. 879–881. Leblanc, M., 1981, The Late Proterozoic ophiolites of Bou Azzer (Morocco): Evidence for Pan-African plate tectonics, in Kröner, A., ed., Precambrian plate tectonics: Amsterdam, Elsevier, p. 435–451. Leblanc, M., and Lancelot, J.R., 1980, Interprétation géodynamique du domaine pan-africain (Précambrien terminal) de l’Anti-Atlas (Maroc) à partir de données géologiques et géochronologiques: Canadian Journal of Earth Sciences, v. 17, p. 142–155. Leblanc, M., and Moussine-Pouchkine, A., 1994, Sedimentary and volcanic evolution of a Neoproterozoic continental margin (Bleida, Anti-Atlas, Morocco): Precambrian Research, v. 70, p. 25–44, doi: 10.1016/03019268(94)90019-1. Leterrier, J., Maury, R.C., Thonon, P., Girard, D., and Marchal, M., 1982, Clinopyroxene composition as a method of identification of the magmatic affinities of paleo-volcanic series: Earth and Planetary Science Letters, v. 59, p. 139–154, doi: 10.1016/0012-821X(82)90122-4. Levresse, G., 2001, Contribution à l’établissement d’un modèle génétique des gisements d’Imiter (Ag-Hg), Bou Madine (Pb-Zn-Cu-Ag-Au), Bou Azzer (Co-Ni-As-Au-Ag) dans l’Anti-Atlas marocain [Ph.D. thesis]: France, University of Nancy, 191 p. Maloof, A.C., Schrag, D.P., Crowley, J.L., and Bowring, S.A., 2005, An expanded record of Early Cambrian carbon cycling from the Anti-Atlas Margin, Morocco: Canadian Journal of Earth Sciences, v. 42, p. 2195– 2216, doi: 10.1139/e05-062. Mifdal, A., and Peucat, J.-J., 1985, Datations U/Pb and Rb/Sr du volcanisme acide de l’Anti-Atlas marocain et du socle sous-jacent dans la région de Ouarzazate. Apport au problème de la limite Précambrien-Cambrien: Sciences Géologiques, Strasbourg, v. 38, p. 185–200. Mifdal, A., Peucat, J.-J., and Charlot, R., 1982, Precambrian-Cambrian U-Pb/ zircon, Morocco, in Odin, G.S., ed., Numerical dating in stratigraphy: Chichester, John Wiley and Sons, p. 941–943. Murphy, J.B., Cameron, K., Dostal, J., Keppie, J.D., and Hynes, A.J., 1985, Cambrian volcanism in Nova Scotia, Canada: Canadian Journal of Earth Sciences, v. 22, p. 599–606. Murphy, J.B., Strachan, R.A., Nance, R.D., Parker, K.D., and Fowler, M.B., 2000, Proto-Avalonia: A 1.2–1.0 Ga tectonothermal event and constraints for the evolution of Rodinia: Geology, v. 28, p. 1071–1074, doi: 10.1130/0091-7613(2000)028<1071:PAAGTE>2.3.CO;2. Nance, R.D., Murphy, J.B., and Keppie, J.D., 2002, A Cordilleran model for the evolution of Avalonia: Tectonophysics, v. 352, p. 11–31, doi: 10.1016/ S0040-1951(02)00187-7. Neltner, L., 1938, Etudes géologiques dans le Sud marocain (Haut-Atlas et Anti-Atlas): Notes et Mémoires, Service Géologique du Maroc, v. 42, p. 1–298. Ouali, H., Briand, B., Bouchardon, J.-L., and El Maâtaoui, M., 2000. Mise en évidence d’un volcanisme alcalin intraplaque d’âge Acadien dans la Meseta nord-occidentale (Maroc): Comptes Rendus Académie des Sciences, Paris, Sciences de la terre et des planètes, v. 330, p. 611–616. Ouali, H., Briand, B., Bouchardon, J.-L., and Capiez, P., 2003, Le volcanisme cambrien du Maroc central: Implications géodynamiques: Comptes Rendus Géoscience, v. 335, p. 425–433, doi: 10.1016/S1631-0713(03)00064-6. Pearce, J.A., 1982, Trace element characteristics of lavas from destructive plate boundaries, in Thorpe, R.S., ed., Andesites: Chichester, John Wiley and Sons, p. 525–548. Piqué, A., 2003, Evidence for an important extensional event during the Latest Proterozoic and Earliest Paleozoic in Morocco: Comptes Rendus Géoscience, v. 335, p. 865–868, doi: 10.1016/j.crte.2003.08.005.
59
Piqué, A., O’Brien, S., King, A.F., Schrenk, P.E., Skehan, J.W., and Hon, R., 1990, La marge nord-occidentale du Paléo-Gondwana (Maroc occidental et zones orientales des Appalaches); Rifting au Précambrien terminal et au Paléozoïque inférieur, et compression hercynienne et alléghanienne au Paléozoïque supérieur: Comptes Rendus Académie des Sciences, Paris, v. 310, série II, p. 411–416. Piqué, A., Bouabdelli, M., and Darboux, J.R., 1995, Le rift cambrien du Maroc occidental: Comptes Rendus Académie des Sciences, Paris, v. 320, série IIa, p. 1017–1024. Piqué, A., Bouabdelli, M., Soulaïmani, A., Youbi, N., and Iliani, M., 1999, Les conglomérats du P III (Néoprotérozoïque supérieur) de l’Anti Atlas (Sud du Maroc): Molasses panafricaines, ou marqueurs d’un rifting fini-Protérozoïque? Comptes Rendus Académie des Sciences, Paris, Sciences de la terre et des planètes, v. 328, p. 409–414. Proust, F., 1973, Étude stratigraphique, pétrographique et structurale du Bloc oriental du Massif ancien du Haut Atlas (Maroc): Notes et Mémoires, Service Géologique du Maroc, v. 34, p. 15–53. Proust, F., Petit, J.-P., and Tapponnier, P., 1977, L’accident du Tizi n’Test et le rôle des décrochements dans la tectonique du Haut Atlas occidental (Maroc): Bulletin de la Société Géologique de France, v. 19, no. 7, p. 541–551. Samson, S.D., Inglis, J.D., D’Lemos, R.S., Admou, H., Blichert-Toft, J., and Hefferan, K., 2004, Geochronological, geochemical, and Nd-Hf isotopic constraints on the origin of Neoproterozoic plagiogranites in the Tasriwine ophiolite, Anti-Atlas orogen, Morocco: Precambrian Research, v. 135, p. 133–147, doi: 10.1016/j.precamres.2004.08.003. Saquaque, A., Admou, H., Karson, J.A., Hefferan, K., and Reuber, I., 1989, Precambrian accretionary tectonics in the Bou Azzer-El Graara region, Anti-Atlas, Morocco: Geology, v. 17, p. 1107–1110, doi: 10.1130/00917613(1989)017<1107:PATITB>2.3.CO;2. Saquaque, A., Benharref, M., Abia, H., Mrini, Z., Reuber, I., and Karson, J.A., 1992, Evidence for a Panafrican volcanic arc and wrench fault tectonics in the Jbel Saghro, Anti-Atlas, Morocco: Geologische Rundschau, v. 81, p. 297–309, doi: 10.1007/BF01764536. Sdzuy, K., 1978, The Precambrian-Cambrian boundary beds in Morocco (preliminary report): Geological Magazine, v. 115, p. 83–94. Seddiki, A., Remaci-Benaouda, N., Cottin, J.-Y., Moine, B.N., Ménot, R.-P., and Perrache, C., 2004, The volcano-sedimentary Boukaïs inlier (southwestern Algeria): Evidence for lithospheric thinning during the Late Neoproterozoic: Journal of African Earth Sciences, v. 39, p. 257–266, doi: 10.1016/j.jafrearsci.2004.07.042. Seifert, K., and Brunotte, D., 1996, Geochemistry of serpentinized mantle peridotite from Site 897 in the Iberia abyssal plain, in Whitmarsh, R.B. et al., eds, Proceedings of ODP, scientific results: College Station, Texas, Ocean Drilling Program, v. 149, p. 413–424. Soulaimani, A., Le Corre, Cl., and Farazdaq, R., 1997, Déformation hercynienne et relation socle/couverture dans le domaine du Bas-Drâa (Anti-Atlas occidental, Maroc): Journal of African Earth Sciences, v. 24, p. 271–284, doi: 10.1016/S0899-5362(97)00043-2. Soulaimani, A., Bouabdelli, M., and Piqué, A., 2003, L’extension continentale au Néo-Protérozoïque supérieur-Cambrien inférieur dans l’Anti-Atlas (Maroc): Bulletin de la Société Géologique de France, v. 174, p. 83–92, doi: 10.2113/174.1.83. Soulaimani, A., Essaifi, A., Youbi, N., and Hafid, A., 2004, Les marqueurs structuraux et magmatiques de l’extension crustale au Protérozoïque terminal-Cambrien basal autour du massif de Kerdous (Anti-Atlas occidental, Maroc): Comptes Rendus Géoscience, v. 336, p. 1433–1441, doi: 10.1016/j.crte.2004.03.020. Soulaimani, A., Jaffal, M., Maacha, L., Kchikah, A., Najine, A., and Saidi, A., 2006, Modélisation magnétique de la suture ophiolitique de Bou AzzerEl Graara (Anti-Atlas central, Maroc). Implications sur la reconstitution géodynamique panafricaine: Comptes Rendus Géoscience, v. 338, p. 153–160, doi: 10.1016/j.crte.2005.10.001. Sun, S.S., and McDonough, W.F., 1989, Chemical and isotopic systematics of oceanic basalts: Implications for mantle composition and processes, in Saunders, A.D., and Norry, M.J., eds, Magmatism in the ocean basins: Geological Society of London, Special Publication, v. 42, p. 313–345. Taylor, S.R., and McLennan, S.M., 1985, The continental crust: Its composition and evolution: Oxford, Blackwell Scientific Publications, 312 p. Thomas, R.J., Chevallier, L.P., Gresse, P.G., Harmer, R.E., Eglington, B.M., Armstrong, R.A., de Beer, C.H., Martini, J.E.J., de Kock, G.S., Macey, P.H., and Ingram, B.A., 2002, Precambrian evolution of the Sirwa
60
Pouclet et al.
Window, Anti-Atlas Orogen, Morocco: Precambrian Research, v. 118, p. 1–57, doi: 10.1016/S0301-9268(02)00075-X. Thomas, R.J., Fekkak, A., Ennih, N., Errami, E., Loughlin, S.C., Gresse, P.G., Chevallier, L.P., and Liégeois, J.-P., 2004, A new lithostratigraphic framework for the Anti-Atlas orogen, Morocco: Journal of African Earth Sciences, v. 39, p. 217–226, doi: 10.1016/j.jafrearsci.2004.07.046. Thorpe, R.S., Francis, P.W., and O’Callaghan, L., 1984, Relative roles of source composition, fractional crystallization and crustal contamination in the petrogenesis of Andean volcanic rocks: Philosophical Transactions of the Royal Society, London, v. A310, p. 675–692. Tucker, R.D., and McKerrow, W.S., 1995, Early Paleozoic chronology: A review in light of new U-Pb zircon ages from Newfoundland and Britain: Canadian Journal of Earth Sciences, v. 32, p. 368–379. Walsh, G.J., Aleinikoff, J.N., Benziane, F., Yazidi, A., and Armstrong, T.R., 2002, U-Pb zircon geochronology of the Paloproterozoic Tagragra de Tata inlier and its Neoproterozoic cover, western Anti-Atlas, Morocco: Precambrian Research, v. 117, p. 1–20, doi: 10.1016/S0301-9268(02)00044-X. Yilmaz, Y., Güner, Y., and Şaroğlu, F., 1998, Geology of the quaternary volcanic centres of the east Anatolia: Journal of Volcanology and
Geothermal Research, v. 85, p. 173–210, doi: 10.1016/S0377-0273 (98)00055-9. Youbi, N., 1998, Le volcanisme “post-collisionel”: Un magmatisme intraplaque relié à des panaches mantelliques. Etude volcanologique et géochimique. Exemples d’application dans le Néoprotérozoïque terminal de l’AntiAtlas et le Permien du Maroc [Ph.D. thesis]: Morocco, University of Marrakech, 737 p. Zahour, G., 2001, Le Néoprotérozoïque terminal de la boutonnière de Toubkal (Haut-Atlas occidental) et de Siroua (Anti-Atlas central): Un exemple de volcanisme intraplaque continental associé à un volcanisme calco-alcalin post-collisionel [Ph.D. thesis]: Casablanca, Morocco, University Hassan II-Mohammedia, 266 p. Zahour, G., Chabanne, A., Youbi, N., and Cabanis, B., 1999, Le volcanisme du Néoprotérozoïque terminal de Tircht (massif ancien du Haut-Atlas occidental, Maroc): Exemples de dômes volcaniques liés à des failles volcanotectoniques héritées du socle protérozoïque: Notes et Mémoires, Service Géologique du Maroc, v. 408, p. 229–236. MANUSCRIPT ACCEPTED BY THE SOCIETY 3 OCTOBER 2006
Printed in the USA
Geological Society of America Special Paper 423 2007
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) Ulf Linnemann* Staatliche Naturhistorische Sammlungen Dresden, Museum für Mineralogie und Geologie, Königsbrücker Landstrasse 159, D-01109 Dresden, Germany Axel Gerdes Institut für Geowissenschaften, Mineralogie, Johann Wolfgang Goethe-Universität Frankfurt am Main, Altenhöferallee 1, D-60438 Frankfurt am Main, Germany Kerstin Drost Staatliche Naturhistorische Sammlungen Dresden, Museum für Mineralogie und Geologie, Königsbrücker Landstrasse 159, D-01109 Dresden, Germany Bernd Buschmann TU Bergakademie Freiberg, Geologisches Institut, Bernhard-von-Cotta-Strasse 2, D-09599 Freiberg, Germany
This paper is dedicated to Jean-Jacques Chauvel (1935–2004)
ABSTRACT Sediment provenances and magmatic events of Late Neoproterozoic (Ediacaran) and Cambro-Ordovician rock complexes from the Saxo-Thuringian zone are constrained by new laser ablation inductively coupled plasma mass spectrometry (LAICP-MS) U-Pb dating of detrital zircons from five sandstones and magmatic zircons from an ignimbrite and one tuffite. These geochronological results in combination with the analysis of the plate-tectonic setting constrained from field observations, sedimentological and geochemical data, and trends of the basin development are used to reconstruct Cadomian orogenic processes during the Late Neoproterozoic and the earliest Cambrian. A continuum between Cadomian orogenesis and the opening of the Rheic Ocean in the Cambro-Ordovician is supported by the data set. *E-mail:
[email protected]. Linnemann, U., Gerdes, A., Drost, K., Buschmann, B., 2007, 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, 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. 61–96, doi: 10.1130/2007.2423(03). For permission to copy, contact
[email protected]. ©2007 Geological Society of America. All rights reserved.
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Linnemann et al. In our model, the early stage of the Cadomian evolution is characterized by a Cordilleran-type continental magmatic arc, which was established at the periphery of the West African craton between ca. 650 and 600 Ma. Subsequently, at ca. 590–560 Ma, a back-arc basin was formed behind the Cadomian magmatic arc. The back-arc basin was closed between ca. 545 and 540 Ma, leading to the development of a short-lived Cadomian retroarc basin. Subsequently, a mid-oceanic ridge was subducted underneath the Cadomian orogen. Slab break-off of the subducted oceanic plate resulted in increased heat flow, leading to voluminous magmatic and anatectic events that culminated at ca. 540 Ma. Oblique incision of the oceanic ridge into the continent caused the formation of rift basins during the Lower to Middle Cambrian. This process continued from the Middle to Upper Cambrian, finally caused the opening of the Rheic Ocean in the Lower Ordovician. Keywords: Peri-Gondwana, Cadomian orogeny, Bohemian Massif, Saxo-Thuringian zone, Cadomia, Avalonia, Rheic Ocean, U-Pb zircon dating, provenance
INTRODUCTION The Cadomian orogeny comprises a series of complex sedimentary, magmatic, and tectonometamorphic events that spanned the mid-Neoproterozoic (ca. 650 Ma) to the earliest Cambrian (ca. 540 Ma). Rock units formed by the Cadomian orogeny are commonly referred to collectively as Cadomian basement. Owing to similar contemporaneous orogenic processes in the Avalonian microplate, the collective term Avalonian-Cadomian orogeny has also been used in the modern literature. Peri-Gondwanan terranes, microcontinents, and crustal units in central, western, and eastern Europe and in north Africa are affected by the Cadomian orogeny. Related orogenic events known as the Avalonian orogeny, are known from the Appalachians (eastern United States and Atlantic Canada) and from the non-Laurentian part of Ireland and the British Isles. Baltica escaped Avalonian-Cadomian orogenic activity, although late Precambrian orogenic events of “Cadomian affinity” have been recognized in the Urals and the Timanides on the periphery of Baltica (Roberts and Siedlecka 2002; Glasmacher et al., 2004). The Cadomian orogeny was first defined in the North Armorican Massif in France on the basis of the unconformity that separates deformed Precambrian rock units from their Early Paleozoic (Cambro-Ordovician) overstep sequence. In central and western Europe this unconformity is commonly referred to as the Cadomian unconformity. The youngest metasedimentary rocks affected by Cadomian deformation may be earliest Cambrian in age, and many geologists assume that the final stages of Cadomian orogenesis were spatially diachronous, lasting from the latest Neoproterozoic to the earliest Cambrian. From this viewpoint the term Cadomian basement includes Neoproterozoic (Ediacaran) to Early Cambrian sedimentary, igneous, and metamorphic complexes, although the stratigraphic range of the rocks involved changes from region to region. The Cadomian unconformity was first described from Rocreux near Caen (Normandy) by Bunel (1835), although it is often attributed to Dufrenoy (1838), who published the first
drawing. The wider geographic extent of the unconformity in central Brittany was recognized by Dufrenoy (1838) and Barrois (1899). The first illustration of the unconformity from Brittany was published by Kerforne (1901). Cadomus and Cadomum are old Latin terms for the modern city of Caen and are the source of the name of the orogeny. The term discordance cadomienne was first used by Bertrand (1921). The type locality of the Cadomian unconformity is located on the northern edge of the village of Jacob Mesnil (Rocreux), close to Bretteville sur Laize near Caen (Normandy). The best illustration of the unconformity at Rocreux was published by Graindor (1957). In some publications, the term Pan-African orogeny is used in the same sense as the Cadomian orogeny, because both events were related to the Gondwana supercontinent in the late Precambrian and occurred at more or less the same time. The main difference between the two orogenic events is their position within the configuration of the Gondwana supercontinent in Neoproterozoic time. The crustal units affected by the Pan-African orogeny are located between the cratons that assembled Gondwana and, in most cases, reflect continent-continent collision (see compilation of Windley, 1995). In contrast, the Cadomian orogen, or alternatively, the Avalonian-Cadomian orogenic belt, was a peripheral orogen at the edge of the Gondwanan supercontinent and is characterized by orogenic processes similar to those of the present-day Andes and Cordilleran chains of the Americas and western Pacific (Murphy and Nance, 1991; Nance and Murphy, 1994; Buschmann, 1995; Linnemann et al., 2000, 2004; Nance et al., 2002). On the basis of provenance studies based on U-Pb ages of detrital zircon grains from sedimentary rocks and inherited zircons in igneous rocks, in combination with Nd-Sr-Pb isotope analyses and paleomagnetic and paleobiogeographic data, most geologists accept that the largest part of the Cadomian basement of central and western Europe was formed at the periphery of the West African craton of the Gondwana supercontinent (e.g., Linnemann et al., 2004; Murphy et al., 2004). Remnants of old cratonic basement are represented only by the Icartian basement
The continuum between Cadomian orogenesis and opening of the Rheic Ocean
Saxo-Thuringian zone (Cadomia) and the Rhenohercynian zone (East Avalonia) during the Late Devonian to Early Carboniferous (Oncken, 1997; Kroner et al., 2003, this volume; Zeh et al., 2003, 2005). To the south and southeast the Bohemian Massif is overthrust by Meso- and Cenozoic rocks of the Alps and the Carpathians. Additional components of the Bohemian Massif are the Teplá-Barrandian unit and the Moravo-Silesian zone. The oldest units of the Bohemian Massif are remnants of Paleo- to Mesoproterozoic cratonic basement slivers, such as the Dobra gneiss (1.38 Ga) and the Svetlik gneiss (2.1–2.05 Ga). The Bohemian Massif is the most prominent inlier of basement rocks in central Europe (Fig. 1) and records a complex Neoproterozoic to Paleozoic history that includes the Cadomian and Variscan orogenies. Some rock units (e.g., the Erzgebirge Mountains) locally experienced ultra-high pressure metamorphic conditions during the Variscan orogeny with the formation of diamond-bearing rocks (Massonne, 1998). Most marginal rock units and inliers, however, were essentially less affected by the Variscan tectonometamorphic overprint. These rock units comprise Neoproterozoic to Paleozoic successions at the northern margin of the Saxo-Thuringian zone and the TepláBarrandian unit (Fig. 1). The Bohemian Massif has been traditionally interpreted to be part of the Armorica microplate as defined by Van der Voo (1979). However, more recent studies (Tait et al., 1997; McKerrow et al., 2000) have assumed that the Armorican microplate was not a coherent block, but comprises several units. Thus, Tait et al. (1997) suggested the term Armorican terrane assemblage, which includes Neoproterozoic and Paleozoic basement units exposed in northern and southern France and in central Europe.
us Iapet
e Sutur
TH OR
Z SX
GEOLOGIC SETTING AM
ltic
a
Cadomi a
BM
S TBU MZ
M
r ne F Alpi
ont
ne Li
The Bohemian Massif forms the central part of the European Variscides and is subdivided into two principal zones, the Saxo-Thuringian zone to the north and Moldanubian zone to the south (Kossmat, 1927). In addition, the marginal Moravo-Silesian zone rims the Bohemian Massif in its eastern part. To the northwest, the massif is bordered by the Mid-German Crystalline zone, which is assumed to represent an important Variscan suture zone, perhaps the Rheic suture (Kroner et al., 2003, this volume; Zeh and Wunderlich, 2003; Zeh et al., 2003, 2005; Linnemann et al., 2004). The latter was closed by oblique collision between the
Ba
ia Avalon Rheic Suture
Rheic Sutu
re
Bohemian Massif
Su tu re
t is qu rn To eyr se is Te
(2.06–2.01 Ga) of the Armorican Massif, and the Svetlik (2.1– 2.05 Ga) and Dobra gneisses (1.38 Ga) of the Bohemian Massif. These Meso- to Paleoproterozoic gneiss complexes coupled with abundant Archean to Paleoproterozoic detrital zircon grains in Neoproterozoic sediments indicate that most of the Cadomian “basement” developed on thinned older cratonic crust and that the Neoproterozoic to Cambrian siliciclastic sediments result from eroded older basement slices. Nance et al. (2002) proposed a Cordilleran model for the evolution of the Neoproterozoic to Cambro-Ordovician rock complexes in the Avalonian part of the “Avalonian-Cadomian orogenic belt.” Avalonia drifted off as a separate microcontinent during the late Cambrian. They suggested that the formation and separation of Avalonia was controlled by a plate-tectonic evolution similar to that presently observed at the western margin of the North American plate in Baja California. Over the past 30 Ma, this area has been affected by terrane accretion, subduction-related processes, ridge-trench collision, and rifting processes. As shown in this article, these processes may also account for the sedimentological and magmatic evolution observed in the Neoproterozoic-Paleozoic basement complexes of the Saxo-Thuringian zone, which lies at the northeastern periphery of the Bohemian Massif. Parts of this zone were less affected by the Variscan orogeny and, thus, forms an ideal area to study sedimentological and magmatic events that occurred during the Neoproterozoic and Cambro-Ordovician. In this article, we present new laser ablation inductively coupled plasma mass spectrometry (LA-ICP-MS) U-Pb data for detrital zircon grains to constrain the provenance of selected sediments of the Saxo-Thuringian zone. In addition, data from zircons in well-defined igneous rocks provide new time markers for the area’s magmatic evolution. These geochronological data, in conjunction with field relations and published results, are used to define the plate-tectonic setting(s) of Neoproterozoic to Cambro-Ordovician rock units of the Saxo-Thuringian zone. Finally, the combined data set is used to present a tentative model for the evolution of the Cadomian basement of the Saxo-Thuringian zone, which starts with the formation of a marginal Cadomian orogen between ca. 650 and 540 Ma and ends with the opening of the Rheic Ocean in the Cambro-Ordovician.
63
FCM
700 km
Figure 1. Location of prominent basement rocks, sutures, continents, and terranes pertinent to this study. AM—Armorican Massif; BM—Bohemian Massif; FCM—French Massif Central; M—Moravo-Silesian zone; MZ—Moldanubian zone; S—Sudetes; SXZ—Saxo-Thuringian zone; TBU—Teplá-Barrandian unit.
64
Linnemann et al.
For the late Neoproterozoic to Cambro-Ordovician basement rocks relevant to Cadomian and post-Cadomian orogenic processes, we use the term Cadomia (sensu Nance and Murphy, 1996), which comprises Cadomian basement rocks (with intercalated older cratonic blocks) of the Armorican Massif, the French Massif Central, and the Bohemian Massif with the exception of the Brunovistulian of the Moravo-Silesian zone (Fig. 1). Based on new geochronological results, the Cadomian basement of the Bohemian Massif can be subdivided into Avaloniantype and Cadomian-type units sensu Murphy et al. (2004). Avalonian-type units contain detrital zircon grains of Mesoproterozoic age, which are assumed to have formed in a juvenile crust between 1.3 and 1.0 Ga. In contrast, Cadomian-type rock units show few or no zircons with ages in the range 1.7–.75 Ga and are dominated by detritus and inherited zircons derived from the West African craton (2.05 Ga and older; Murphy et al. 2004). To date, all U-Pb zircon provenance studies indicate that the Neoproterozoic to Paleozoic sediments of the Saxo-Thuringian zone, the Teplá-Barrandian unit, and the Moldanubian zone have a west African provenance (Linnemann et al. 2000, 2004; Gehmlich, 2003; Tichomirowa, 2003; Drost et al., 2004). Thus, they belong to Cadomia sensu Murphy et al. (2004). However, protolith ages of ca. 2.1 Ma indicate that the Svetlik granite gneiss of the Bohemian Massif was emplaced during the Paleoproterozoic (Wendt et al., 1993, 1994). From its geological position it seems likely that this gneiss represents part of the Eburnian basement derived from the West African craton. However, Finger et al. (2000) demonstrated that the Brunovistulian unit in the Moravo-Silesian zone of the Bohemian Massif (Fig. 1) shows strong affinities with Avalonia. Rocks from the Brunovistulian unit are assumed to have been derived from the recycled margin of the Amazonian craton, as suggested by U-Pb ages of inherited zircon grains (Friedl et al., 2000). The 1.38-Ma Dobra gneiss is assumed to represent a cratonic inlier that also belongs to the Avalonian part of the Bohemian Massif (Gebauer and Friedl, 1994; Friedl et al., 2004). The available ages suggest that the Bohemian Massif is divided by a suture (likely the Rheic suture) into Avalonian and Cadomian parts (Fig. 1). The Cadomian part comprises the Saxo-Thuringian and Moldanubian zones, the Teplá-Barrandian unit, and the Sudetes, whereas the Avalonian part includes the Brunovistulian unit of the Moravo-Silesian zone. The former “Rheic” suture is probably hidden under and/or incorporated into the Variscan fault-and-thrust belt between the Moravo-Silesian and the Moldanubian zones (Fig. 1). Saxo-Thuringian Zone The Saxo-Thuringian zone forms the northeastern part of the Bohemian Massif. It consists of Cadomian basement units, which are overlain by a Paleozoic overstep sequence (Fig. 2). The parautochthonous part of the Saxo-Thuringian zone forms a northeast– east-trending fold-and-thrust belt, which consists of the Schwarzburg antiform, the North Saxon antiform, the Berga antiform, and
the Lausitz antiform, and the Torgau-Doberlug and ZiegenrückTeuschnitz synclines. In addition, the Saxo-Thuringian zone is transected by the northwest–southeast-trending Elbe zone and the Franconian line (Fig. 2). In this study we use the neutral word antiform instead of the traditional term anticline, because none of the tectonostratigraphic units are typical anticlines. For example, the Lausitz antiform is a tilted horst block. The Saxo-Thuringian zone in this article is subdivided into an internal and external domain, which show significant differences with respect to their Cadomian basement evolution and Paleozoic overstep sequences. The external domain is composed of the Cadomian volcanosedimentary units of the Rothstein Formation in the Torgau-Doberlug syncline and the Altenfeld Formation in the northwestern part of the Schwarzburg antiform (Figs. 2 and 3). Both are characterized by rock units containing thick layers of massive black chert (Fig. 4A) and are assumed to have originated in a back-arc setting (Buschmann, 1995; Linnemann et al., 2000). These sediments are dominated by dark-gray to black distal turbidites composed of an intercalation of graywacke and mudstone bedsets. All known sedimentological and geochemical data point to an origin for the Rothstein Formation in the center of a back-arc basin developed on thinned continental crust (Buschmann, 1995). Owing to its similar spatial position in the Saxo-Thuringian zone and its similarity in lithology and geochemistry, we assign the Altenfeld Formation to the same plate-tectonic setting. Zircon data suggest deposition of the Rothstein and Altenfeld formations at ca. 570–565 Ma (Linnemann et al., 2000; Buschmann et al., 2001). The Rothstein Formation is overlain by Lower to Middle Cambrian sediments (Fig. 5), whereas the Altenfeld Formation is covered by Lower Ordovician siliciclastics (Fig. 6). In contrast to the internal domain, ca. 540-Ma magmatism in the external domain is very scarce. Only a single small pre-Variscan granitoid body (the Milchberg granite) crops out in the northwestern part of the Schwarzburg antiform, where it intrudes the Altenfeld Formation. Recent U-Pb zircon datings place that granite to the base of the Ordovician (489 ± 6 Ma; U. Linnemann and A. Gerdes, unpublished data; Fig. 3). Another important component of the external domain is the Vesser complex of Middle to Upper Cambrian age. This unique complex is characterized by rocks related to the formation of the oceanic crust (Bankwitz et al., 1992; Kemnitz et al., 2002). The relationship between the external domain and the Mid-German Crystalline zone to the north is unclear because of coverage by Cenozoic sediments. The bounding element of the internal domain is the Blumenau Shear Zone, which divides the Schwarzburg antiform into a northwestern and southeastern part. In our view, the Blumenau Shear Zone continues to the southern border of the Torgau-Doberlug syncline, which is also covered by Cenozoic deposits. The shear zone is a structural feature that likely originated in the Cadomian orogeny during the tectonic change from a back-arc basin to a retroarc basin setting (see below). During the Variscan orogeny, the Blumenau Shear Zone was reactivated as a sinistral shear zone (Heuse et al., 2001).
n
2
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map
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Rhe ic
16°
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Low to high grade ortho- and para-rocks Ordovician, Silurian, and Devonian volcano-sedimentary rock complexes Cambro-Ordovician to Lower Carboniferous volcano-sedimentary rock complexes
Variscan early molasses of HainichenBorna (Upper Viséan) Variscan early molasses of Doberlug (Upper Viséan)
Sudetes
Olistolithes of Cambrian to Devonian rock complexes within a wildflysch matrix Acid to basic metamorphic rocks of the nappe pile remnants of Münchberg and of the Saxon "Zwischengebirge" of Wildenfels and Frankenberg
Variscan wildflysch with large olistolithes
Lower Graptolite Shale and Ockerkalk (Silurian) Carbonates, sandstones, pelites, and diabases (Devonian) Greywackes and pelites (Tournai, Visé) (Variscan flysch)
Drillings Heinersdorf 1 & 2 (Lower to Middle Cambrian)
Bavarian Facies (Variscan Wildflysch and Variscan nappe piles)
H
(Geological map without any strata younger than Lower Carboniferous)
lt
Sutu re
Cadomia
tra-
Av
n In
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t)
Praha
a ni alo
Balt
The Saxo-Thuringian Zone (NE Bohemian Massif)
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e ab
F
Dresden
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us
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15°
ain: dom rnal ian e t x E om Cad rc basin -a in: back oma nal d ian r e t In om Cad rc basin -a retro
14°
Germany
Compiled by Ulf Linnemann and Manfred Schauer (1999) Dresden, April 5, 1999
hr
n ia
e irg
Freiberg Frankenberg
E
b
Rothstein
b nu e a n d ol Zo M
Wildenfels
30
c
6
e - rg lit bi nu ge Hainichen ra G
Cadomian passive margin
5 D
d
13°
Cadomian and Ordovician rocks affected by the dextral Variscan Blumenau shear zone (Schwarzburg antiform)
km
4
G Leipzig
Dessau
Zo ne
Tourmaline granite (Elbe zone) Western Lausitz granitoids
Variscan and post - Variscan igneous rocks (ca. 330 - 300 Ma) Granitoids of the Leipzig area Granitoids of the Elbe zone (Dohna Permo-Carboniferous & Laas granodiorites) granitoids shear zone-related Upper Carboniferous orthogneisses rhyolitoids (Grossenhain gneiss) Upper Carboniferous Granitoids of the major granitoid dikes Schwarzburg antiform
Eastern Lausitz granitoids
Rumburk granite (Lausitz antiform)
Ordovician granitoids (ca. 490 Ma)
Lausitz anatexite
Cadomian igneous rocks (Lower Cambrian, c. 540 - 530 Ma)
Surface outcrops
Rheno-Herzynian zone
Southern phyllite zone (Cambro-Ordovician rock complex)
Vesser complex (Upper Cambrian)
Northern phyllite zone (Cambro-Ordovician rock complex)
Surface outcrops of the Mid-German crystalline zone
Low to high grade ortho- and para-rock complexes of the Mid-German crystalline zone and related rock units
Adjoining areas of the Saxo-Thuringian zone
Phyllites and garnet phyllites (Mid Pressure - Low Temperature unit and Low Pressure - Low Temperature unit)
Gneisses, eclogites and mica schists with major shear zones (High Pressure - High Temperature unit)
"Red" ortho-gneisses, anatexites, migmatites (Mid Pressure - Mid Temperature unit)
Erzgebirge and Fichtelgebirge
High grade basaltic rocks and serpentinites
High grade country rocks of the granulite core
Granulite
Sächsisches Granulitgebirge ("Saxon Granulite massif") (Variscan metamorphic core complex)
65
Figure 2. Geological map of the Saxo-Thuringian zone in the northeastern part of the Bohemian Massif, showing units of Lower Carboniferous and older ages and the distribution of rocks that represent the different stages of Cadomian basin development (modified from Linnemann and Schauer, 1999; Linnemann and Romer, 2002). Tectonostratigraphic units: 1—Schwarzburg antiform (southeastern part); 2—Schwarzburg antiform (northwestern part); 3—Berga antiform; 4—North Saxon antiform (Leipzig area); 5—North Saxon antiform (Clanzschwitz area); 6—Torgau-Doberlug syncline; 7—Elbe zone; 8—Lausitz antiform. Neoproterozoic volcano-sedimentary complexes: A—Rothstein Formation; B—Altenfeld Formation; C—Frohnberg Formation; D—Clanzschwitz Group; E—Rödern Group; F—Weesenstein Group; G—Leipzig Formation; H—Lausitz Group. Sample locations (small type in stars): a—sample Pur-1 (Purpurberg quartzite of the Weesenstein Group, Neoproterozoic); b—Wett-1 (microconglomerate of the Lausitz Group, Neoproterozoic); c—Roth-1–1641H/18 (graywacke of the Rothstein Formation, Neoproterozoic); d—Kam-1–1209/1 (sandstone of the Zwethau Formation, Lower Cambrian); e—Lbq-1 (microconglomerate, Langer Berg Formation, Tremadoc, Lower Ordovician).
Carbonates, sandstones, and pelites (Lower to Middle Cambrian) Quartzites and pelites (Skolithos facies) (Collmberg Fm., Hainichen-Otterwisch Fm., Dubrau Fm.) (Lower Ordovician, Tremadoc) Quartzites, shales, sed. iron ore (Ordovician)
3
Plauen
Gera
Cadomian basement (Neoproterozoic) (sedimentary rocks, synsedimentary volcanic rocks) Graywackes, pelites, cherts, volcanic rocks (Altenfeld Fm., Frohnberg Fm., Lausitz Group, Leipzig Formation) Graywackes, pelites, cherts, basalts, andesites (Rothstein Formation) Graywackes, pelites, quartzites, and basalts (passive margin sequences and tillites of the Clanzschwitz, Rödern, and Weesenstein Gr.) Thuringian facies (Palaeozoic sedimentary rocks and volcanic rocks)
Sa
o- Z
Jena
Saalfeld Schwarzburg
an
rm
Zo
ne
ec 12° ht in ge n
Fl
ne
n
li al
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y Cr
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ia cyn
B ai m do ian sin : C l a m a ain rn o b n te ad -arc om ian sin ia x d m a C k E l g Fr c a o b an ba ern ad arc r i n co u nt C roe I ni t 50° h n a re T o
n:
er
He
Kyffhäuser
e -G
id M
T W hü al rin d g
Ruhla
o-
z
Vesser Complex Langer Berg
51°
e Rh
N
ar
Ehrenberg
H
Ilmenau
11°
Lausitz granitoid complex
10° 52°
The continuum between Cadomian orogenesis and opening of the Rheic Ocean
U M
3
Vesser 502+/-2 complex Milchberg granite
L
489+/-6
5
530+/-8
2
483+/-3
487+/-6
2
555+/-9
4
533+/-4 541+/-7
2
2
Cadomian basement senu stricto
Neoproterozoic (Ediacaran)
542 Ma relative paleo-position: North
566+/-10
dated tuff
dated detrital zircon
Lausitz Group
Leipzig Formation
1
4
629+/-4
2
569+/-2
2
589+/-9
Proximate part of the retro-arc basin against the magmatic arc
1 Altenfeld Formation
2
485+/-6
Laas granodiorite
531+/-7
Dohna granodiorite
537+/-7
2
2
relative paleo-position: South
Frohnberg Formation 577+/-3
Roth-1
Rothstein Formation
dated granitoid pebble
Wett-1
*
551+/-8
2
Tourmaline granite
2
Glasbach & Laubach granites
539+/-6
Elbe zone 486+/-4
2
2
Lausitz granitoid complex
Schildau massif
605+/-4 570+/-4
Blambach rhyolite
Rumburk granite 490+/-3
2
Cadomian Orogeny sensu stricto (deformation close to the Prec./Camb.Boundary, magmatic event at c. 540 Ma)
Lower Ordovician
no Ordovician sediments
488 Ma
Cambrian
North North Saxon Linnemann et al. Schwarzburg Saxon antiform antiform antiform Lausitz (Leipzig) antiform (Clanzschwitz) (SE-Part) 479+/-2
Schwarzburg TorgauDoberlug antiform syncline (NW-Part)
66
Upper section* (~100m): remnant basin of the retro-arc against the passive margin & cratonic crust
577+/-10
2
568+/-4
2 2
Pur-1
Weesenstein Group
Clanzschwitz Group
Lower section: Passive margin of part of the distal the back-arc basin on retro-arc basin thinned cratonic against the crust and remnants magmatic arc of an older Cadomian magmatic arc Cadomian retro-arc basin
Strike-slip and spreading zones in the back-arc basin on thinned continental crust
Cadomian back-arc basin (S)
Cadomian back-arc basin (N)
External domain
Internal domain
Parautochthonous part of the Saxo-Thuringian zone dated subvolcanic intrusion
1
2
3
4
5
6
7
8
9
10
Figure 3. Generalized lithostratigraphic profiles of parautochthonous units of the Saxo-Thuringian zone, with published geochronological data of the Cadomian basement and its Cambro-Ordovician overstep sequence. Circles designated “Roth-1,” “Wett-1,” and “Pur-1” indicate position of samples studied in this article. 1—Cambro-Ordovician rift-related igneous rocks; 2—Cadomian granitoids of the ca. 540-Ma magmatic event; 3—Lower Ordovician siliciclastic sediments; 4—Late Neoproterozoic debris flows and glaciomarine tillites; 5—igneous rocks and metasediments of the Upper Cambrian Vesser complex (predominantly mafic rocks); 6—Neoproterozoic hydrothermal black cherts; 7—Lower to Middle Cambrian sediments; 8—conglomerates, quartzites, and quartzitic shales of the Purpurberg quartzite (Weesenstein Group) and its equivalent in the Clanzschwitz Group; 9—graywackes and mudstones; 10—predominantly mudstones. Sources of geochronological data (numbered circles): 1—SHRIMP U-Pb (Buschmann et al., 2001); 2—thermal ionization mass spectrometer (TIMS) Pb-Pb (Linnemann et al., 2000); 3—TIMS U-Pb (Kemnitz et al., 2002); 4—SHRIMP U-Pb (Linnemann et al., 2004); 5—laser ablation inductively coupled plasma mass spectrometry (LA-3CP MS) U-Pb (Linnemann and Gerdes, unpublished data).
The continuum between Cadomian orogenesis and opening of the Rheic Ocean
A
C
67
B
D
E
Figure 4. Photographs of Cadomian basement rocks of the Saxo-Thuringian zone (Bohemian Massif). (A) Bedded black cherts of the “Rothstein Rock” near Bad Liebenwerda, which are interpreted to have formed close to the spreading center of the Cadomian back-arc basin (Neoproterozoic, ca. 570 Ma, Rothstein Formation, Cadomian basement of the Torgau-Doberlug syncline). (B) Steeply dipping bed of the Purpurberg quartzite with internal cross bedding and impressions of wave ripples (left freestanding bedding plane of block above hammer). The quartzite occurs in the lower part of the Weesenstein Group, which is interpreted to represent the passive margin of the Cadomian back-arc basin (Neoproterozoic, ca. 570 Ma, Weesenstein Group, Purpurberg near Niederseidewitz, Cadomian basement of the Elbe zone). (C) Stretched granitoid pebble in a mudstone matrix derived from an eroded magmatic arc during the formation of the Cadomian back-arc basin (upper part of the Weesenstein Group). The pebbly mudstone facies of the Weesenstein Group is interpreted to be in part glaciomarine (Neoproterozoic, ca. 570 Ma, Weesenstein Group, 100 m east of the Weesenstein railway station, Cadomian basement of the Elbe zone). (D) Microconglomerate of the Lausitz Group containing clasts of cherts and felsic rocks. This sediment demonstrates the redeposition of black cherts and arc volcanics during the formation of the Cadomian retroarc basin (Neoproterozoic, ca. 570–545 Ma, Lausitz group, Petershain near Kamenz, Cadomian basement of the Lausitz anticline). (E) Water escape structure at top of the a-interval of a graywacke turbidite of the Lausitz Group, demonstrating rapid sedimentation in the Cadomian retroarc basin (Neoproterozoic, ca. 570–545 Ma, Lausitz Group, Wetterberg quarry near Ebersbach, Cadomian basement of the Lausitz anticline).
?
?
?
?
?
?
D IV (> 41 m)
?
D I (> 126 m)
LS 1/63I (> 575 m)
?
Kam-1
Torgau Member (> 500 m)
Delitzsch Formation Tröbitz Formation
?
Rosenfeld Mbr. (> 300 m)
Zwethau Formation
The internal domain of the Saxo-Thuringian zone contains Cadomian rock units from two different depositional settings (Fig. 3). Voluminous Cadomian plutons intruded at ca. 540– 530 Ma, and a thick and widely distributed Ordovician overstep sequence (Fig. 6) also distinguish the internal domain. Cambrian deposits are restricted to the Heinersdorf 1 and 2 drill holes in the Berga antiform and to large olistolithes in a Lower Carboniferous wild flysch matrix in the Görlitz synform. The first group of Neoproterozoic sedimentary units in the internal domain comprises passive margin sequences characterized by highly mature quartzites, sandstones, and quartz-rich shales deposited in a shallow marine environment (Linnemann, 1991). The most prominent deposit of this type is represented by the Purpurberg quartzite in the lower part of the Weesenstein Group (Fig. 4B). Facies analysis of the Purpurberg quartzite has shown that its deposition was caused by an extreme drop of sea level that is interpreted to be of glacioeustatic origin (Linnemann, 1991). The stratigraphic equivalent of the Purpurberg quartzite also occurs in the Clanzschwitz Group (North Saxon antiform). Other parts of the Weesenstein and Clanzschwitz groups are likewise passive margin deposits but comprise quartz-rich mud- and siltstones. In the upper part of the two groups diamictites and layers with isolated pebbles (Fig. 4C) may be glaciomarine in origin (Linnemann and Romer, 2002). These passive margin deposits are situated in the North Saxon antiform and the Elbe zone (Fig. 2). Based on the spatial arrangement of the passive margin deposits in the Saxo-Thuringian zone, we assign these units to the passive margin of the same Cadomian back-arc basin in which the Rothstein and the Altenfeld formations were deposited. Detrital zircons and dated pebbles point to a minimum age of sedimentation of ca. 570 Ma for the passive margin units (Linnemann et al., 2000; Fig. 3). The second group of Neoproterozoic sedimentary units in the internal domain is represented by the Lausitz Group (Lausitz antiform), the Leipzig Formation (North Saxon antiform), and the Frohnberg Formation (southeastern part of the Schwarzburg antiform). All three units are characterized by monotonous, flyschlike sections of proximal to distal dark-gray to black colored turbidites composed of graywacke and mudstone couplets (Fig. 4E). Seismites indicate an active tectonic setting during basin formation. Intercalations of conglomerates contain fragments of black cherts (Fig. 4D) and other debris from older Neoproterozoic sediments and igneous rocks. The pervasive occurrence of fragments of black chert in both the graywackes and the conglomerates suggests that deposits from the Cadomian back-arc basin of the external zone became eroded, recycled, and redeposited in the Cadomian retroarc basin, remnants of which are represented by the Lausitz Group and the Leipzig and the Frohnberg formations. Sensitive high-resolution ion microprobe (SHRIMP) U-Pb dating of detrital zircon grains in the Lausitz Group and the Frohnberg Formation indicates that they are younger than 555 ± 9 Ma and 551 ± 8 Ma, respectively (Linnemann et al., 2004). Because of the occurrence of debris derived from the back-arc basin, the presence of distinct sedimentary features (see below), and their
Middle Cambrian
Linnemann et al.
Lower Cambrian
68
claystone
mafic volcanics
siltstone
dolostone
sandstone siliciclastic debris flows
limestone ?
no record
Figure 5. Generalized lithostratigraphic profile of Cambrian sediments of the Torgau-Doberlug syncline with stratigraphic position of sample Kam-1–1209/1. The Lower Cambrian profile is documented in numerous drill cores, whereas the Middle Cambrian profile is derived from reference profiles (boreholes D I, LS 1/63, D IV). The Lower to Middle Cambrian sediments in the column overlie the Rothstein Formation (Neoproterozoic). Mbr—member. From Buschmann et al. (2006).
5
11
6
12
NW-part of the Schwarzburg antiform
Vesser complex
Cadomian basement
Tremadoc
Middle ore horizon Lower ore horizon 3000
Phycodes quartzite
2000 Phycodes shale
Upper Magnetite quartzite Dachschiefer 1000 Fm.
1000 m
“Phycodes Dachschiefer” Lower Magnetite quartzite
Ves-1 Gabbro 502+/-2 Ma *2
Oberer Frauenbachquartzite
Bärentiegel porphyroid (479 +/-5 Ma*1)
FrauenbachWechsellagerung
interbedding of shale and quartzite beds
Unterer Frauenbachquartzite
Lower Frauenbach quartzite
Upper Frauenbach quartzite
Goldisthal Fm.
Dacitic pyroclastite 508+/-2 Ma *2
Frauenbach Gr.
Vesser complex
Neuwerk Fm.
Volcanic unit
Lbq-1
Ordovician
Quartzite unit
Langer Berg Fm.
Rollkopf Formation
Tremadoc Upper Cambrian
Vesser Group
Middle Cambr.
Hundsrück G.
Gillersdorf Fm.
gap
Griffelschiefer Fm.
Quarzitbank Member
10
Banded Lederschiefer “Kalkbank” (limestone layer) Upper ore horizon
Schmiedefeld Fm.
Phycodenquarzit Fm.
4
gap and/or condensed sedimentation
69
Quarzitplatten Mbr.
9
Phycodes Group
3
schiefer Fm.
Phycodenschiefer Formation Lauschenstein Member Rosen- BreitenGöritzberg Member berg Mbr. berg lower upper Mbr.
2
Arenig Ashgill Gräfenthal G.
SE-part of the Schwarzburg antiform 7 The continuum between Cadomian orogenesis and opening of the Rheic Ocean glaciomarine diamictite (m) of the Sahara glaciation Leder8 = Lederschiefer
1
gap
?
no outcrop
0
dark shales
0
gap Cadomian basement: Milchberg granite Altenfeld Group (Neoproterozoic) black cherts
gap
KArc-1
Conglomeratic tuffite (”Konglomeratische Arkose”) and yellow tuffites Blambach rhyolite (487+/-5 Ma*1) Cadomian basement: topmost Neoproteroz. Quartzite Glasbach granite (538+/-4 Ma*1) Frohnberg Group (Neoprot.)
Figure 6. Lithostratigraphic profiles of the Middle and Upper Cambrian and Ordovician rocks of the Vesser complex and the northwestern and southeastern parts of the Schwarzburg antiform. Ellipses Ves-1, Lbq-1, and KArc-1 indicate approximate position of samples studied in this article. 1—Upper Cambrian Rollkopf Formation of the Vesser complex: predominantly mafic subvolcanic rocks (tholeiitic dolerites and gabbros) and minor subalkaline basalts, dacitic tuffs, rhyolitic ignimbrites, granites, and graphitic metasediments; 2—Upper Cambrian Neuwerk Formation of the Vesser complex: interbedded tholeiitic basaltic, dacitic to trachyandesitic lavas and intermediate to rhyolitic pyroclastics and metasediments; 3—Tremadocian volcanic unit of the Hundsrück Group in the Vesser complex: rhyolitic pyroclastic rocks, and arkoses; 4—rhyolites and porphyroids; 5—Cadomian granites; 6—Neoproterozoic sediments (predominantly graywacke turbidites); 7—conglomerates, microconglomerates, and conglomeratic tuffites; 8—sandstones and quartzites; 9—mudstones and silty shales; 10—shales; 11—diamictite (glaciomarine tillite?); 12—sedimentary iron ores. Fm—formation; G—group; Mbr—member. Sources of geochronological data: *1—TIMS Pb-Pb (Linnemann et al., 2000); *2—TIMS U-Pb (Kemnitz et al., 2002). Modified after Bankwitz et al. (1992), Linnemann (1996), Linnemann and Heuse (2000), and Kemnitz et al. (2002).
70
Linnemann et al.
maximum ages of sedimentation between ca. 555 and 551 Ma, we classify the Lausitz Group and the Leipzig and Frohnberg formations as relicts of a Cadomian retroarc basin. In general, all Neoproterozoic sections within the Cadomian basement of the Saxo-Thuringian zone seem to be rootless because of Variscan stacking of the crust, and an underlying cratonic basement is not known. However, neodymium depletedmantle (NdTDM) model ages for the Late Neoproterozoic sediments range between 1.9 and 1.3 Ga (Linnemann and Romer, 2002), clearly indicating that the source area of the Neoproterozoic sediments was dominated by old cratonic crust. With exception of the Rothstein Formation and maybe the Altenfeld Formation, all Neoproterozoic sedimentary sequences within the Cadomian basement in the Saxo-Thuringian zone were intruded by Early Cambrian post-kinematic granitoid plutons in the interval ca. 540–530 Ma (Linnemann et al., 2000; Gehmlich, 2003; Tichomirowa, 2003). These plutonic suites are composed of granites, syeno- and monzogranites, granodiorites, and tonalites (Hammer, 1996), whereas granodiorites dominate in most plutons. The Cadomian basement of the Saxo-Thuringian zone is overlain, usually unconformably, by Lower Paleozoic sediments. Transgression and the development of Lower to Middle Cambrian overstep sequences—including the deposition of conglomerates, carbonates, siliciclastics, and red beds, with a depositional gap in the lowermost Cambrian (ca. 540–530 Ma)— characterize the first post-Cadomian sedimentary sequence. A second widely distributed gap in sedimentation occurred in the Upper Cambrian (ca. 500–490 Ma), although the Vesser complex is composed of mid- to Upper Cambrian magmatic rocks and metasediments related to an oceanic setting (Bankwitz et al., 1992; Kemnitz et al., 2002). Special features—such as the occurrence of a Cadomian unconformity; peri-Gondwanan Cambro-Ordovician faunas; glaciomarine diamictites of the Hirnantian glaciation in the uppermost Ordovician; and the absence of any Salinic, Acadian, and Caledonian orogenic influences—paleogeographically link the Saxo-Thuringian zone to Gondwana in the Neoproterozoic and Lower Paleozoic (Linnemann et al., 2000, 2004). SAMPLES AND METHODS For provenance studies, detrital zircons were collected from three Neoproterozoic siliciclastic sedimentary rock units, which were deposited in three distinct settings of Cadomian basin development. Sample Pur-1 was taken from the Purpurberg quartzite of the Weesenstein Group in the Elbe zone. This sediment was deposited in a passive continental margin setting of the Cadomian back-arc basin distal from the arc. Sample Roth-1 is a graywacke of the Rothstein Formation in the Torgau-Doberlug syncline taken from the drill hole WisBaW 1641H/80 near the city of Herzberg. Sediments of the Rothstein Formation were deposited in the Cadomian back-arc basin proximal to the arc on the opposite side to that of the passive margin.
The third Neoproterozoic sample is a chert-bearing microconglomerate (Wett-1) of the Lausitz Group from the Lausitz antiform. This sample was collected from the Wetterberg quarry near the village of Ebersbach. The Lausitz Group is dominated by graywacke turbidites with intercalations of microconglomerates deposited in the Cadomian retroarc basin or foreland basin. In addition, a Lower Cambrian sandstone (Kam-1) and an Ordovician microconglomerate (Lbq-1) were sampled from the Saxo-Thuringian zone. These samples represent Cambro-Ordovician shelf sediments, which overlie the Cadomian basement. Kam-1 was taken from a drill core of the Zwethau Formation in the Torgau-Doberlug syncline. The sample was collected from drill hole WisBaW 1209/78 near the village of Falkenberg and is representative of the Lower Cambrian overstep sequence overlying the deformed Cadomian sediments of the Rothstein Formation. The Lower to Middle Cambrian sediments of this formation were deposited in an asymmetric rift basin. Lbq-1 is a Lower Ordovician microconglomerate sampled from the Langer Berg Formation close to the village of Willmersdorf in the northwestern part of the Schwarzburg antiform. The Langer Berg Formation is a section of highly mature quartzites and conglomerates typical of the widely distributed Lower Ordovician shallow marine sedimentation of the Gondwanan realm. The Lower Ordovician overstep sequence in the Saxo-Thuringian zone was deposited in a rifted shelf basin in a passive margin setting. To set additional lithostratigraphic time markers for the Cambrian and Ordovician sedimentation, an Upper Cambrian ignimbrite (Ves-1) and a Lower Ordovician tuffite (KArc-1) were sampled. Sample Ves-1 was taken from a rhyolitic ignimbrite from the Vesser complex (Fig. 6). Sample KArc-1, a pebble-bearing rhyolitic tuffite, was collected in the valley of the Blambach close to Sitzendorf, from the base of the >3000-m-thick Ordovician sedimentary succession exposed in the southeastern part of the Schwarzburg antiform. This pyroclastic sediment is referred to in traditional German literature as “Konglomeratische Arkose” (= conglomeratic arkose). Additional information concerning the lithostratigraphy and coordinates of the sample locations is given in Tables 1, 2, and 3. Zircon concentrates were separated at the Museum für Mineralogie und Geologie (Staatliche Naturhistorische Sammlungen Dresden). Fresh samples were crushed in a jaw crusher and sieved for the fraction 63–400 μm. Density separation of this fraction by a heavy liquid was realized using sodium heteropolytungstate in water (“LST fast float”) and followed by magnetic separation of the extracted heavy minerals in a Frantz isodynamic separator. Final selection of the zircon grains for U-Pb dating was achieved by hand-picking under a binocular microscope. Zircon grains of all grain sizes and morphological types were selected, mounted in resin blocks, and polished to half their thickness. Zircons were analyzed for U, Th, and Pb isotopes by LAICP-MS techniques at the Institute of Geosciences, Johann Wolfgang Goethe-University Frankfurt, using a Thermo-Finnigan Element II™ sector field ICP-MS coupled to a New Wave™ UP-213 ultraviolet laser system. A teardrop-shaped, low-volume laser cell
Pb* (cps)
U (ppm)
†
Pb (ppm)
†
Th U
†
Pb U
238
206
1σ (%)
Pb U
235
207
1σ (%) 206
207
Pb Pb
1σ %
Rho**
Pb U
235
207
±2σ (Ma)
Ages Pb U
238
206
±2σ (Ma) 206
207
Pb Pb
±2σ (Ma)
P-L10-1 P-L10-2 P-L10-3 P-L10-4 P-L10-5 P-L10-6 P-L10-7 P-L10-8 P-L10-9 P-L10-10 P-L10-11 P-L10-12 P-L10-13 P-L10-14 P-L10-15 P-L10-16 P-L10-17 P-L10-18 P-L10-19 P-L10-20 P-L10-21 P-L10-22 P-L10-23 P-L10-24 P-L10-25 P-L10-26 P-L10-27 P-L10-28 P-L10-29 P-L10-30 P-L10-31 P-L10-32 P-L10-33 P-L10-35 P-L10-36 P-L10-37 P-L10-38 P-L10-39 P-L10-40 P-L10-41
5820 26148 6324 1024 3723 18901 1033 958 18098 39418 1114 9860 15049 1739 1010 27258 4213 1704 5850 1934 5516 4759 7390 1702 57171 24872 1243 1118 2173 1679 1506 3245 988 898 3227 3112 12438 1015 1235 683
160 681 206 231 96 857 73 73 145 405 248 182 222 170 66 174 60 24 162 259 97 62 82 248 371 169 197 221 369 296 60 96 75 174 444 415 267 81 43 123
64 284 99 25 46 197 14 10 101 194 25 21 140 17 8 141 35 11 67 28 47 44 53 27 322 139 21 27 63 48 24 41 13 20 59 52 121 13 18 15
0.31 0.44 1.85 0.72 1.15 0.13 1.40 1.07 0.26 0.47 0.58 1.15 1.01 0.67 0.93 0.51 0.62 0.92 0.68 0.86 0.98 1.38 0.51 0.75 0.82 0.46 0.32 0.56 2.08 1.50 0.68 0.63 1.69 0.43 0.86 0.44 0.43 0.90 1.05 0.47
0.3770 0.3820 0.3091 0.0957 0.3739 0.2281 0.1293 0.1167 0.6011 0.3566 0.0925 0.0925 0.4907 0.0924 0.0981 0.6481 0.4918 0.4004 0.3513 0.0926 0.3409 0.5083 0.5505 0.0970 0.6531 0.6679 0.0919 0.1005 0.0979 0.1033 0.3134 0.3321 0.1011 0.0961 0.1036 0.1064 0.3637 0.0995 0.3301 0.0997
0.9 0.8 1.0 0.9 0.9 1.0 1.3 1.0 1.0 1.0 1.0 1.1 0.9 1.0 1.2 0.8 1.1 1.7 0.9 0.8 0.8 1.1 1.0 1.0 0.9 1.1 1.4 1.6 1.4 1.2 1.7 1.2 1.2 1.7 1.2 1.1 1.2 1.4 1.5 1.2
6.966 6.775 5.406 0.7893 6.433 3.704 2.305 1.041 22.33 13.38 0.7705 0.7486 11.85 0.7566 0.8093 26.72 12.25 7.900 6.098 0.7673 5.642 13.00 14.87 0.7922 26.09 26.36 0.7537 0.8280 0.8156 0.8878 4.730 5.350 0.8390 0.8117 0.8752 0.9252 6.371 0.8480 5.319 0.8389
1.4 1.0 1.3 2.2 1.6 1.1 3.2 2.7 1.3 1.7 2.1 1.7 1.1 2.1 2.4 1.0 1.6 2.5 1.4 2.1 2.7 1.5 1.5 2.0 1.0 1.4 2.5 2.5 1.7 2.0 2.3 1.6 2.6 3.3 1.8 2.1 1.5 2.6 2.5 2.0
0.1340 0.1287 0.1268 0.0598 0.1248 0.1178 0.1293 0.0647 0.2694 0.2721 0.0604 0.0587 0.1751 0.0594 0.0598 0.2991 0.1807 0.1431 0.1259 0.0601 0.1200 0.1856 0.1959 0.0592 0.2898 0.2863 0.0595 0.0597 0.0604 0.0623 0.1094 0.1168 0.0602 0.0613 0.0613 0.0631 0.1270 0.0618 0.1169 0.0610
1.1 0.6 0.9 2.0 1.2 0.6 2.9 2.5 0.9 1.3 1.8 1.4 0.7 1.9 2.1 0.7 1.2 1.8 1.1 1.9 2.6 1.1 1.1 1.8 0.4 0.7 2.0 1.9 1.1 1.5 1.5 1.0 2.3 2.8 1.3 1.8 0.9 2.1 2.0 1.7
0.62 0.81 0.71 0.41 0.61 0.85 0.42 0.38 0.73 0.62 0.49 0.62 0.80 0.47 0.51 0.78 0.66 0.68 0.61 0.38 0.29 0.70 0.65 0.49 0.90 0.84 0.56 0.63 0.79 0.62 0.75 0.75 0.47 0.53 0.66 0.52 0.82 0.57 0.59 0.58
2107 2083 1886 591 2037 1572 1214 725 3198 2707 580 567 2592 572 602 3373 2624 2220 1990 578 1922 2680 2807 592 3350 3360 570 613 606 645 1773 1877 619 603 638 665 2028 624 1872 619
25 17 23 20 27 18 46 29 26 32 18 15 21 19 22 20 30 46 25 19 47 29 28 18 20 27 22 23 16 19 39 27 24 30 17 21 27 24 43 19
2062 2085 1736 589 2047 1325 784 712 3034 1966 571 570 2574 570 603 3221 2578 2171 1941 571 1891 2649 2827 597 3240 3298 567 617 602 634 1758 1849 621 592 635 652 2000 612 1839 612
31 28 29 10 33 23 20 14 47 35 11 12 39 11 14 41 45 64 29 9 26 46 44 11 46 59 15 19 16 15 53 38 14 20 14 14 43 17 47 14
2151 2080 2054 598 2026 1922 2088 765 3302 3318 617 555 2607 581 597 3465 2659 2265 2041 608 1957 2703 2792 576 3416 3398 585 595 618 686 1790 1908 610 648 649 710 2057 667 1909 641
39 20 34 88 44 21 103 107 29 41 78 60 23 81 91 20 40 64 40 84 91 36 37 78 13 23 88 84 46 65 55 37 99 121 58 77 31 90 71 71
Pur-1 (Location: Purpurberg quartzite, Neoproterozoic, Edicarian, Weesentein group, Elbe zone, Purpurberg near Oberseidewitz, Easting: 42 3117, Northing: 56 40750)
Sample
207
Isotopic ratios§
TABLE 1. LASER ABLATION-ICP-MS U, Pb, AND Th DATA OF DETRITAL ZIRCON GRAINS FROM NEOPROTEROZOIC QUARZITES OF THE ELBEZONE, SAXO-THURINGIAN ZONE, BOHEMIAN MASSIF
Continued
96 100 85 99 101 69 38 93 92 59 92 103 99 98 101 93 97 96 95 94 97 98 101 104 95 97 97 104 97 92 98 97 102 91 98 92 97 92 96 96
Conc (%)
The continuum between Cadomian orogenesis and opening of the Rheic Ocean 71
Pb* (cps)
207
1616 5059 1543 1201 1277 12501 1460 30561 1147 3728 1068 10912 9971 2136 1453 8669 5387 13890 999 21756 1192 1300 10484 989 7933 5692 1396 3073 5608 1230 4643 41824 3440 4675
312 65 87 75 97 121 133 178 150 91 28 303 284 62 51 170 110 286 98 413 242 86 251 216 72 107 230 73 68 85 168 386 533 105
U† (ppm) 33 40 19 10 12 31 29 168 20 50 13 108 98 24 7 80 44 118 10 157 25 29 124 20 56 47 25 33 40 9 86 247 61 49
Pb† (ppm) 0.23 0.37 3.06 0.64 0.58 0.45 0.92 0.59 0.60 1.22 0.77 0.45 0.45 0.69 0.59 0.68 0.37 0.46 0.38 0.98 0.56 0.07 1.12 0.36 1.06 0.39 0.79 0.81 0.54 0.04 0.34 0.43 0.55 0.71
Th† U 0.0935 0.4832 0.1042 0.0992 0.1017 0.1459 0.1626 0.6604 0.1049 0.3744 0.3359 0.3266 0.3167 0.3370 0.1296 0.4035 0.3636 0.3685 0.0952 0.3100 0.0957 0.3501 0.3659 0.0904 0.5863 0.3871 0.0981 0.3714 0.5189 0.1101 0.4687 0.5523 0.1040 0.3911
Pb 238 U
206
1.1 1.1 1.5 1.5 1.4 1.8 1.2 1.2 1.4 1.1 1.3 1.3 1.4 1.5 1.4 1.4 1.5 1.4 1.4 3.3 1.3 1.7 1.4 1.5 1.3 1.8 1.5 1.3 1.4 2.0 1.3 1.6 1.5 1.4
1σ (%) 0.7746 11.40 0.8843 0.8400 0.8696 1.426 1.642 25.14 0.8823 6.417 5.504 5.258 5.072 5.496 1.179 7.459 6.518 6.955 0.7844 6.783 0.7881 5.634 6.155 0.7364 17.15 8.612 0.8144 6.529 13.46 0.9314 11.01 16.19 0.8929 7.344
Pb 235 U
207
1.7 1.5 2.9 3.0 2.4 2.2 1.8 1.5 2.4 1.6 1.8 1.6 1.7 2.1 2.8 1.6 2.7 1.6 2.9 3.4 2.4 2.6 1.5 2.8 1.5 2.4 2.7 2.0 1.8 8.4 1.6 1.6 1.8 1.8
1σ (%) 0.0601 0.1711 0.0616 0.0614 0.0620 0.0709 0.0733 0.2760 0.0610 0.1243 0.1189 0.1168 0.1161 0.1183 0.0660 0.1341 0.1300 0.1369 0.0597 0.1587 0.0597 0.1167 0.1220 0.0591 0.2122 0.1614 0.0602 0.1275 0.1882 0.0614 0.1704 0.2127 0.0623 0.1362
Pb 206 Pb
207
1.4 0.9 2.5 2.6 2.0 1.3 1.4 0.9 2.0 1.1 1.2 0.9 1.0 1.4 2.4 0.9 2.2 0.7 2.6 0.9 2.0 2.0 0.7 2.4 0.8 1.6 2.3 1.5 1.0 8.2 0.9 0.4 1.1 1.1
1σ % 0.63 0.77 0.51 0.49 0.56 0.81 0.63 0.81 0.57 0.71 0.73 0.81 0.82 0.72 0.51 0.85 0.57 0.90 0.48 0.96 0.53 0.65 0.88 0.52 0.85 0.75 0.54 0.66 0.82 0.23 0.81 0.96 0.81 0.80
Rho**
582 2557 643 619 635 900 987 3314 642 2035 1901 1862 1832 1900 791 2168 2048 2106 588 2084 590 1921 1998 560 2943 2298 605 2050 2713 668 2524 2888 648 2154
Pb 235 U
207
16 28 28 28 23 27 23 29 23 29 31 27 30 36 31 30 48 28 26 62 22 45 27 24 30 44 25 35 33 84 30 31 17 32
±2σ (Ma) 576 2541 639 609 624 878 971 3269 643 2050 1867 1822 1774 1872 786 2185 1999 2022 586 1741 589 1935 2010 558 2974 2109 603 2036 2695 673 2478 2835 638 2128
±2σ (Ma) 12 48 18 17 16 30 21 61 17 40 43 41 44 48 21 52 53 49 16 101 15 56 47 16 63 64 17 46 63 25 54 72 18 52
Ages Pb 238 U
206
Pb Pb
607 2569 659 655 675 954 1021 3341 639 2019 1939 1907 1898 1931 805 2152 2098 2188 594 2442 594 1906 1986 570 2922 2470 611 2064 2726 652 2561 2926 683 2179
206
207
58 32 106 113 87 53 58 27 85 40 44 33 35 51 101 30 78 24 111 32 88 71 26 105 26 53 98 53 33 353 32 14 46 37
±2σ (Ma)
W-s1-u1 W-s1-u2 W-s1-u3 W-s1-u4 W-s1-u5 W-s1-u6 W-s1-u7
3134 6397 6264 5911 11710 12003 95089
69 159 157 100 255 281 108
8 16 13 11 28 30 58
0.41 0.44 0.18 0.62 0.42 0.53 0.13
0.1051 0.0956 0.0878 0.0959 0.1051 0.0997 0.4985
0.7 0.7 0.6 0.7 0.6 0.6 0.6
0.8834 0.7771 0.7057 0.7851 0.8774 0.8162 13.78
1.3 1.2 1.3 1.6 1.0 1.0 0.9
0.0610 0.0589 0.0583 0.0594 0.0605 0.0594 0.2004
1.1 1.0 1.2 1.5 0.9 0.8 0.6
0.55 0.54 0.46 0.44 0.57 0.58 0.71
643 584 542 588 640 606 2734
12 11 11 14 10 9 16
644 589 542 590 644 613 2607
9 8 6 8 7 7 26
638 565 542 580 622 581 2830
46 46 51 63 37 36 20
Wett-1 (Location: microconglomerate, Neoproterozoic, Ediacarian, Lausitz group, Lausitz antiform, Wetterberg near Ebersbach, Easting: 40 5284, Northing: 56 80226)
P-L10-42 P-L10-43 P-L10-44 P-L10-45 P-L10-46 P-L10-47 P-L10-48 P-L10-49 P-L10-50 P-L10-51 P-L10-52 P-L10-53 P-L10-54 P-L10-55 P-L10-56 P-L10-57 P-L10-58 P-L10-59 P-L10-60 P-L10-61 P-L10-62 P-L10-63 P-L10-64 P-L10-65 P-L10-66 P-L10-67 P-L10-68 P-L10-69 P-L10-70 P-L10-71 P-L10-74 P-L10-76 P-L10-77 P-L10-78
Pur-1 (continued)
Sample
Isotopic ratios§
TABLE 1. LASER ABLATION-ICP-MS U, Pb, AND Th DATA OF DETRITAL ZIRCON GRAINS FROM NEOPROTEROZOIC QUARZITES OF THE ELBEZONE, SAXO-THURINGIAN ZONE, BOHEMIAN MASSIF (continued)
Continued
101 104 100 102 104 105 92
95 99 97 93 93 92 95 98 101 102 96 96 93 97 98 102 95 92 99 71 99 102 101 98 102 85 99 99 99 103 97 97 93 98
Conc (%)
72 Linnemann et al.
Pb* (cps)
207
W-s1-u7c W-s1-u8 W-s1-u9 W-s1-u10 W-s1-u11 W-s1-u12 W-s1-u13 W-s1-u14 W-s1-u15 W-s1-u16 W-s1-u17 W-s1-u18 W-s1-u19 W-s1-u20 W-s1-u21 W-s1-u22 W-s1-u23 W-s1-u24 W-s1-u25 W-s1-u26 W-s1-u27 W-s1-u28 W-s1-u29 W-s1-u30 W-s2-u31 W-s2-u32 W-s2-u33 W-s2-u34 W-s2-u35 W-s2-u36 W-s2-u37 W-s2-u38 W-s2-u39 W-s2-u40 W-s2-u41 W-s2-u42 W-s2-u43 W-s2-u44 W-s2-u45 W-s2-u46 W-s2-u47
195089 17962 8833 10565 26352 34944 70514 18605 19303 289166 19542 40604 4567 5922 12664 35868 51721 25772 6473 41036 14934 3528 29962 77679 12942 14820 260309 27117 19538 5354 37659 4056 5770 25160 61665 11239 15790 6575 35955 20520 8530
Wett-1 (continued)
Sample
309 376 199 224 94 120 255 411 263 398 80 61 92 143 312 182 157 69 148 1061 267 82 81 74 290 51 762 405 397 63 105 74 140 82 82 138 371 148 147 52 201
U† (ppm)
214 46 21 24 37 44 96 43 45 229 28 35 12 15 32 53 66 26 21 98 35 11 34 58 35 21 307 64 47 11 47 9 14 36 56 23 38 18 52 26 20
Pb† (ppm)
0.35 0.86 0.46 0.38 0.79 0.35 0.65 0.56 0.43 0.24 0.67 0.72 0.64 0.52 0.49 0.33 0.58 0.45 1.60 0.27 0.74 1.58 0.42 1.12 1.24 0.95 0.30 0.49 0.70 0.59 0.84 0.32 0.56 1.03 0.81 0.44 0.51 0.83 0.56 0.92 0.67
Th† U
0.6099 0.1013 0.0966 0.1012 0.3309 0.3450 0.3275 0.0960 0.1627 0.5290 0.2945 0.4717 0.1169 0.0961 0.0956 0.2744 0.3673 0.3332 0.0987 0.0923 0.1001 0.0956 0.3860 0.5687 0.0923 0.3328 0.3799 0.1498 0.1056 0.1613 0.3629 0.1151 0.0924 0.3506 0.5392 0.1537 0.0960 0.1038 0.3194 0.3920 0.0884
Pb 238 U
206
1.6 0.7 0.6 0.7 0.8 0.7 0.6 0.8 0.8 0.6 1.0 0.6 0.7 0.6 0.6 1.2 0.7 1.2 0.8 0.8 0.7 0.6 0.8 0.6 1.1 0.9 1.0 0.8 0.9 0.8 0.8 0.9 0.8 0.8 0.9 1.1 0.9 0.9 0.9 0.9 0.9
1σ (%)
20.97 0.8543 0.7851 0.8557 5.228 5.519 5.089 0.7817 1.608 13.40 4.530 11.46 1.0112 0.7856 0.7783 4.111 6.174 5.627 0.8126 0.7495 0.8509 0.7955 6.859 18.01 0.7625 5.237 6.542 1.414 0.8925 1.5905 6.354 0.9976 0.7630 5.626 14.03 1.456 0.7963 0.8667 4.725 7.314 0.7206
Pb 235 U
207
1.8 1.0 1.0 1.1 1.0 0.9 0.8 1.0 1.1 0.7 1.2 0.8 1.4 1.2 1.1 1.3 0.9 1.9 1.2 0.9 1.0 1.5 1.0 0.7 1.3 1.3 1.0 1.2 1.2 1.2 0.9 1.5 1.3 1.0 1.1 1.3 1.1 1.4 1.1 1.2 1.4
1σ (%)
Isotopic ratios§
0.2494 0.0611 0.0589 0.0613 0.1146 0.1160 0.1127 0.0591 0.0717 0.1838 0.1116 0.1762 0.0627 0.0593 0.0590 0.1087 0.1219 0.1225 0.0597 0.0589 0.0617 0.0603 0.1289 0.2297 0.0599 0.1141 0.1249 0.0685 0.0613 0.0715 0.1270 0.0628 0.0599 0.1164 0.1886 0.0687 0.0602 0.0606 0.1073 0.1353 0.0591
Pb 206 Pb
207
0.9 0.6 0.8 0.9 0.6 0.5 0.5 0.6 0.7 0.3 0.6 0.5 1.3 1.0 0.9 0.5 0.5 1.4 0.9 0.4 0.7 1.4 0.6 0.4 0.8 0.9 0.4 0.9 0.8 0.9 0.4 1.2 1.0 0.6 0.6 0.8 0.7 1.1 0.7 0.7 1.1
1σ %
0.88 0.75 0.57 0.62 0.81 0.84 0.80 0.78 0.78 0.86 0.84 0.75 0.48 0.53 0.54 0.91 0.84 0.66 0.66 0.91 0.74 0.42 0.80 0.85 0.79 0.72 0.93 0.63 0.75 0.68 0.88 0.57 0.63 0.80 0.85 0.80 0.79 0.62 0.75 0.79 0.61
Rho**
3137 627 588 628 1857 1904 1834 586 973 2708 1736 2561 709 589 585 1657 2001 1920 604 568 625 594 2093 2990 575 1859 2052 895 648 966 2026 703 576 1920 2751 912 595 634 1772 2151 551
Pb 235 U
207
36 9 9 10 16 15 14 9 13 13 20 14 15 11 10 21 16 32 11 8 9 14 18 14 12 22 18 15 12 15 16 15 12 18 20 16 10 13 19 21 12
±2σ (Ma)
Ages
3070 622 595 621 1843 1910 1826 591 972 2737 1664 2491 713 591 589 1563 2017 1854 607 569 615 589 2104 2903 569 1852 2076 900 647 964 1996 702 570 1937 2780 922 591 637 1787 2132 546
Pb 238 U
206
78 9 6 8 25 25 20 9 15 26 29 24 9 7 7 33 26 40 9 9 9 7 28 29 11 30 34 13 11 15 27 11 9 28 40 18 10 11 27 33 9
±2σ (Ma) Pb Pb
3181 644 565 651 1873 1896 1843 570 976 2687 1825 2618 699 579 568 1777 1984 1992 594 562 663 616 2083 3050 600 1866 2027 883 650 972 2057 703 600 1901 2730 890 609 624 1754 2168 571
206
207
27 27 35 37 20 17 17 28 27 11 23 17 54 44 41 20 18 50 39 17 29 60 21 12 36 32 13 39 35 36 15 52 45 22 18 33 30 48 27 24 48
±2σ (Ma)
TABLE 1. LASER ABLATION-ICP-MS U, Pb, AND Th DATA OF DETRITAL ZIRCON GRAINS FROM NEOPROTEROZOIC QUARZITES OF THE ELBEZONE, SAXO-THURINGIAN ZONE, BOHEMIAN MASSIF (continued)
Continued
96 97 105 95 98 101 99 104 100 102 91 95 102 102 104 88 102 93 102 101 93 96 101 95 95 99 102 102 100 99 97 100 95 102 102 104 97 102 102 98 96
Conc (%)
The continuum between Cadomian orogenesis and opening of the Rheic Ocean 73
Pb* (cps)
15685 314944 3838 6277 689884 6300 3523 18814 4091 34192 11118 5488 8843
370 914 79 124 1239 152 70 393 50 868 283 136 199
U (ppm)
†
37 323 8 14 498 14 8 35 5 81 27 13 24
Pb (ppm)
†
0.54 0.12 0.52 0.67 0.16 0.35 0.51 0.14 0.57 0.35 0.31 0.28 0.86
Th U
†
Pb U
0.0936 0.3489 0.0875 0.1029 0.3707 0.0878 0.1048 0.0932 0.0881 0.0923 0.0937 0.0953 0.1027
238
206
1.0 1.1 0.8 0.9 1.4 0.8 1.0 1.0 1.1 1.0 0.8 0.9 1.0
1σ (%) Pb U
0.7674 6.117 0.7031 0.8740 9.602 0.7039 0.8761 0.7601 0.7016 0.7466 0.7631 0.7819 0.8485
235
207
1.2 1.2 1.6 2.6 1.8 1.3 1.8 1.3 1.8 1.2 1.1 1.7 1.3
1σ (%)
Isotopic ratios§ Pb Pb
0.0595 0.1272 0.0583 0.0616 0.1879 0.0582 0.0607 0.0591 0.0578 0.0587 0.0591 0.0595 0.0599
206
207
0.7 0.3 1.4 2.4 1.0 1.1 1.4 0.9 1.5 0.6 0.8 1.4 0.9
1σ %
0.83 0.96 0.51 0.34 0.81 0.57 0.59 0.76 0.58 0.86 0.74 0.56 0.72
Rho** Pb U
578 1993 541 638 2397 541 639 574 540 566 576 587 624
235
207
10 20 13 24 33 11 17 12 15 10 10 15 17
±2σ (Ma)
Ages Pb U
577 1929 541 631 2033 542 642 575 544 569 577 587 630
238
206
11 37 8 11 50 8 13 11 11 11 9 10 12
±2σ (Ma) Pb Pb
585 2059 540 661 2724 536 627 572 521 555 570 585 602
206
207
29 12 60 103 34 48 62 37 66 26 33 59 40
±2σ (Ma)
99 94 100 95 75 101 102 100 104 103 101 100 105
Conc (%)
Rot-1 Rot-2 Rot-3 Rot-4 Rot-5 Rot-6 Rot-7 Rot-8 Rot-9 Rot-10 Rot-11 Rot-12 Rot-13 Rot-14 Rot-15 Rot-16 Rot-17 Rot-18 Rot-19 Rot-20 Rot-21 Rot-22 Rot-23 Rot-24 Rot-25
2913 13391 4375 5690 2711 3461 7087 4626 7383 2220 157667 5806 4588 3948 8184 1896 4317 4317 4677 17606 5410 16527 37002 18123 62569
57 338 112 151 64 97 195 121 207 49 593 162 122 106 39 32 127 127 133 466 88 437 1059 89 221
6 40 12 15 6 9 22 15 24 7 233 16 14 12 14 5 13 13 14 54 12 50 107 29 83
0.55 0.93 0.66 0.43 0.31 0.37 0.83 1.70 1.52 1.01 0.60 0.58 0.90 1.03 0.58 0.57 0.61 0.61 0.46 1.09 1.03 0.83 0.55 0.36 0.19
0.0917 0.1011 0.0996 0.0967 0.0926 0.0951 0.0986 0.0965 0.0929 0.1157 0.3531 0.0912 0.0994 0.0922 0.3073 0.1304 0.0942 0.0943 0.0991 0.0958 0.1045 0.0992 0.0932 0.3136 0.3674
1.0 1.0 1.0 1.1 1.2 1.2 1.0 1.0 1.0 1.4 1.1 1.0 1.1 1.0 1.0 1.1 1.2 1.2 1.0 1.1 1.0 1.1 1.1 1.0 1.0
0.7416 0.8465 0.8402 0.8132 0.7511 0.7870 0.8162 0.7967 0.7643 1.001 5.936 0.7370 0.8227 0.7500 4.623 1.172 0.7699 0.7735 0.8165 0.8061 0.8766 0.8320 0.7792 4.709 6.546
1.6 1.3 1.3 1.4 2.0 2.1 1.4 1.5 1.4 1.7 1.2 1.3 1.6 1.8 1.5 1.6 2.1 2.2 1.4 1.4 1.6 1.3 1.2 1.2 1.3
0.0587 0.0607 0.0612 0.0610 0.0588 0.0600 0.0601 0.0599 0.0597 0.0627 0.1219 0.0586 0.0600 0.0590 0.1091 0.0652 0.0593 0.0595 0.0598 0.0610 0.0609 0.0608 0.0606 0.1089 0.1292
1.2 0.8 0.9 0.9 1.6 1.7 1.0 1.1 1.0 1.0 0.4 0.7 1.2 1.5 1.0 1.2 1.7 1.8 0.9 0.7 1.2 0.7 0.5 0.7 0.8
0.65 0.79 0.75 0.78 0.60 0.59 0.71 0.69 0.72 0.82 0.93 0.81 0.69 0.58 0.72 0.67 0.58 0.55 0.74 0.83 0.64 0.86 0.92 0.83 0.79
563 623 619 604 569 589 606 595 577 704 1966 561 610 568 1753 787 580 582 606 600 639 615 585 1769 2052
14 12 13 13 18 18 13 13 13 17 20 11 15 16 25 18 18 20 13 12 15 12 11 21 23
565 621 612 595 571 586 606 594 573 706 1949 563 611 568 1727 790 580 581 609 590 641 610 575 1758 2017
11 12 12 12 13 14 11 11 11 18 36 11 13 11 32 16 13 13 12 13 12 13 13 32 36
555 630 645 639 561 603 605 599 592 699 1985 552 605 567 1785 780 577 584 595 639 634 634 626 1781 2088
52 35 38 37 71 72 43 46 43 41 15 33 50 63 37 50 73 80 41 32 52 28 21 25 28
Continued
102 99 95 93 102 97 100 99 97 101 98 102 101 100 97 101 101 99 102 92 101 96 92 99 97
Roth-1 (Location: graywacke, Neoproterozic, Ediacarian, Rothstein Formation, Torgau-Doberlug syncline, drill WisBaW 1641H/80 near Herzberg, depth of sampled core: 507.0 m, Easting: 38 3904, Northing: 57 27419)
W-s2-u48 W-s2-u49 W-s2-u50 W-s2-u51 W-s2-u52 W-s2-u53 W-s2-u54 W-s2-u55 W-s2-u56 W-s2-u57 W-s2-u58 W-s2-u59 W-s2-u60
Wett-1 (continued)
Sample
207
TABLE 1. LASER ABLATION-ICP-MS U, Pb, AND Th DATA OF DETRITAL ZIRCON GRAINS FROM NEOPROTEROZOIC QUARZITES OF THE ELBEZONE, SAXO-THURINGIAN ZONE, BOHEMIAN MASSIF (continued)
74 Linnemann et al.
Pb* (cps)
207
3860 19366 2559 10046 8780 83961 390248 3281 7044 1477 2656 9091 4856 2962 11390 6613 57107 11895 8343 69989 3745 3756 3580 5581 2471 4638 2468 86195 8981 16769 10693 49118 1797 18536 210159 283975
85 59 59 219 231 312 462 54 199 42 68 272 142 89 306 87 114 355 35 155 118 100 45 170 45 155 40 284 271 60 302 159 40 438 311 488
U† (ppm)
11 31 7 28 25 136 283 7 21 4 8 26 15 11 42 17 60 39 14 79 13 11 5 18 5 16 4 118 31 24 36 69 5 42 168 193
Pb† (ppm)
1.18 1.34 0.76 0.56 0.60 0.73 0.15 0.41 0.52 0.47 0.87 0.46 0.63 1.14 1.36 0.99 0.39 0.86 0.72 0.34 0.80 0.58 0.84 0.69 0.97 0.90 0.47 0.06 0.84 0.42 1.01 0.58 0.63 0.30 0.51 0.26
Th† U
0.1113 0.3847 0.1048 0.1161 0.0991 0.3770 0.5542 0.1218 0.0988 0.0948 0.1051 0.0925 0.0934 0.0965 0.1055 0.1605 0.4688 0.0965 0.3517 0.4540 0.0945 0.1061 0.0974 0.0946 0.0946 0.0895 0.1028 0.4132 0.1013 0.3725 0.1001 0.3834 0.1225 0.0951 0.4614 0.3449
Pb 238 U
206
1.2 1.1 1.4 1.0 1.1 0.9 0.9 1.3 1.0 0.9 0.9 1.1 1.0 0.9 1.3 1.0 1.0 0.9 1.0 1.0 1.2 1.1 0.9 1.1 1.0 1.0 1.2 1.0 1.1 1.1 1.0 0.9 1.0 1.2 3.0 1.8
1σ (%)
0.9605 7.280 0.8798 1.022 0.8242 6.572 17.585 1.0882 0.8155 0.7748 0.8946 0.7525 0.7783 0.7982 0.8858 1.595 12.171 0.7985 5.588 10.970 0.7722 0.9046 0.8043 0.7800 0.7763 0.7308 0.8624 7.717 0.8433 6.508 0.8384 7.104 1.079 0.7890 9.874 5.897
Pb 235 U
207
1.7 1.2 2.3 1.3 1.4 1.0 1.0 1.8 1.4 2.3 1.4 1.4 1.5 1.4 1.6 1.5 1.2 1.3 1.4 1.2 1.6 1.5 1.9 1.4 1.7 1.3 1.5 1.1 1.4 1.4 1.4 1.0 1.5 1.4 7.4 2.4
1σ (%)
0.0626 0.1372 0.0609 0.0638 0.0603 0.1264 0.2302 0.0648 0.0599 0.0593 0.0617 0.0590 0.0604 0.0600 0.0609 0.0721 0.1883 0.0600 0.1153 0.1752 0.0592 0.0619 0.0599 0.0598 0.0595 0.0592 0.0608 0.1355 0.0604 0.1267 0.0607 0.1344 0.0639 0.0602 0.1552 0.1240
Pb 206 Pb
207
1.2 0.7 1.8 0.9 0.9 0.5 0.3 1.2 1.0 2.1 1.1 0.9 1.1 1.0 0.9 1.2 0.6 0.9 1.0 0.6 1.1 1.0 1.7 0.9 1.3 0.9 0.9 0.5 0.8 0.9 0.9 0.5 1.0 0.8 6.7 1.6
1σ %
0.71 0.85 0.60 0.76 0.78 0.87 0.96 0.72 0.70 0.40 0.65 0.76 0.66 0.69 0.83 0.65 0.88 0.73 0.72 0.84 0.75 0.76 0.49 0.78 0.60 0.74 0.77 0.90 0.81 0.75 0.73 0.88 0.72 0.83 0.41 0.74
Rho**
684 2146 641 715 610 2056 2967 748 606 582 649 570 585 596 644 968 2618 596 1914 2521 581 654 599 585 583 557 631 2199 621 2047 618 2125 743 591 2423 1961
Pb 235 U
207
17 22 22 14 13 18 19 19 13 20 14 12 13 12 15 19 22 11 24 22 14 14 18 13 15 11 14 19 13 25 13 18 15 13 140 42
±2σ (Ma)
680 2098 643 708 609 2062 2842 741 607 584 644 570 576 594 646 959 2478 594 1943 2413 582 650 599 583 583 552 631 2229 622 2041 615 2092 745 585 2446 1910
±2σ (Ma)
16 38 17 14 13 31 43 18 12 10 11 12 11 11 16 17 43 10 34 40 13 14 11 12 11 10 14 36 13 38 12 32 15 13 124 59
Ages Pb 238 U
206
Pb Pb
694 2193 634 735 616 2049 3053 768 600 578 665 567 619 603 636 989 2727 603 1884 2608 576 669 600 597 586 576 634 2170 616 2053 630 2156 737 611 2404 2015
206
207
51 23 79 36 38 18 9 52 44 91 46 41 47 42 38 47 18 37 35 21 46 41 73 38 58 38 41 16 35 33 40 17 43 34 228 57
±2σ (Ma)
†
Notes: Coordinates are UTM World Geodetic System 84. Conc.—concordance. *Within-run background-corrected mean 207Pb signal in counts per second. U and Pb content and Th/U ratios were calculated relative to GJ-1 and are accurate to ~10%. § Corrected for background, mass bias, laser induced U-Pb fractionation and common Pb (if detectable, see text on analytical method) using the Stacey and Kramers (1975) model Pb composition. 207Pb/235U calculated using 207Pb/206Pb/(238U/206Pb × 1/137.88). Errors are propagated by quadratic addition of within-run errors (1 standard error) and the reproducibility of GJ-1 (1 standard deviation). **Rho is the error correlation defined as err206Pb/238U/err207Pb/235U. See text for details.
Rot-26 Rot-27 Rot-28 Rot-29 Rot-30 Rot-31 Rot-32 Rot-33 Rot-34 Rot-35 Rot-36 Rot-37 Rot-38 Rot-39 Rot-40 Rot-41 Rot-42 Rot-43 Rot-44 Rot-45 Rot-46 Rot-47 Rot-48 Rot-49 Rot-50 Rot-51 Rot-52 Rot-53 Rot-54 Rot-55 Rot-56 Rot-57 Rot-58 Rot-59 Rot-60c Rot-60r
Roth-1 (continued)
Sample
Isotopic ratios§
TABLE 1. LASER ABLATION-ICP-MS U, Pb, AND Th DATA OF DETRITAL ZIRCON GRAINS FROM NEOPROTEROZOIC QUARZITES OF THE ELBEZONE, SAXO-THURINGIAN ZONE, BOHEMIAN MASSIF (continued)
98 96 101 96 99 101 93 96 101 101 97 101 93 99 102 97 91 98 103 93 101 97 100 98 99 96 100 103 101 99 98 97 101 96 102 95
Conc (%)
The continuum between Cadomian orogenesis and opening of the Rheic Ocean 75
Pb* (cps)
207
U† (ppm)
Pb† (ppm)
Th† U Pb 238 U
206
1σ (%)
Pb 235 U
207
1σ (%)
Pb 206 Pb
207
1σ (%)
Rho**
Pb 235 U
207
±2σ (Ma)
Ages Pb 238 U
206
±2σ (Ma) 206
207
Pb Pb
±2σ (Ma) Conc (%)
Continued
Kam-1 (Location: sandstone, Lower Cambrian, Zwethau Formation, Torgau-Doberlung syncline, drill WisBaW 1209/78 near Falkenberg, depth of sample core: 455.5 m, Easting: 37 4090, Northing: 57 18911) Kam-1 2733 40 6 1.48 0.1046 1.1 0.8747 2.2 0.0606 1.9 0.49 638 21 641 13 626 82 102 Kam-2 72050 222 86 0.26 0.3637 1.1 6.4764 1.1 0.1291 0.4 0.92 2043 20 2000 36 2086 16 96 Kam-3 8653 201 30 1.71 0.1018 1.1 0.8569 1.4 0.0610 1.0 0.74 628 14 625 13 641 42 98 Kam-4 6249 118 16 0.78 0.1131 1.0 0.9648 1.7 0.0619 1.4 0.59 686 17 691 13 670 59 103 Kam-5 7664 186 20 0.49 0.1000 1.2 0.8387 1.6 0.0609 1.0 0.78 618 15 614 14 634 42 97 Kam-6 30020 800 74 0.22 0.0943 1.1 0.7658 1.3 0.0589 0.7 0.85 577 12 581 13 564 30 103 Kam-7 5408 134 17 1.27 0.0966 1.1 0.8008 1.5 0.0601 1.0 0.74 597 13 594 12 608 43 98 Kam-8 2415 45 5 0.44 0.1044 1.1 0.8836 2.0 0.0614 1.7 0.54 643 19 640 13 652 71 98 Kam-9 4321 106 12 0.51 0.1017 1.1 0.8476 1.7 0.0604 1.3 0.65 623 16 625 13 619 57 101 Kam-10 3880 80 8 0.49 0.0982 1.1 0.8272 1.7 0.0611 1.3 0.63 612 16 604 12 641 57 94 Kam-11 2940 49 5 0.62 0.0967 1.2 0.8097 1.8 0.0607 1.4 0.62 602 17 595 13 630 62 94 Kam-12 41770 87 49 0.97 0.4415 1.1 9.868 1.2 0.1621 0.5 0.92 2423 22 2357 44 2478 16 95 Kam-13 76368 340 116 0.62 0.3020 1.2 4.639 1.3 0.1114 0.5 0.93 1756 22 1701 36 1823 17 93 Kam-14 8013 188 23 0.96 0.1008 1.1 0.8391 1.5 0.0604 1.0 0.72 619 14 619 13 618 45 100 Kam-15 250631 667 275 0.38 0.3749 1.2 7.074 1.3 0.1369 0.5 0.93 2121 23 2052 43 2188 17 94 Kam-16 14500 405 38 0.41 0.0885 1.0 0.7168 1.2 0.0587 0.6 0.85 549 10 547 11 557 28 98 Kam-17 2843 54 7 0.91 0.1095 1.2 0.9221 2.4 0.0611 2.0 0.52 663 23 670 16 643 87 104 Kam-18 89961 850 142 0.04 0.1742 1.1 2.255 1.5 0.0939 1.0 0.73 1198 21 1035 20 1506 38 69 Kam-19 8018 179 22 0.84 0.1042 1.1 0.8847 1.6 0.0616 1.1 0.71 644 15 639 14 659 48 97 Kam-20 8018 179 22 0.84 0.1042 1.1 0.8865 1.5 0.0617 1.0 0.74 644 14 639 14 665 43 96 Kam-21 7400 193 20 0.72 0.0928 1.1 0.7667 1.4 0.0599 1.0 0.73 578 13 572 12 601 43 95 Kam-22 3342 82 9 0.75 0.0983 1.2 0.8052 1.8 0.0594 1.4 0.63 600 17 605 13 582 61 104 Kam-23 35884 109 54 1.24 0.3742 1.1 6.654 1.2 0.1290 0.6 0.87 2067 22 2049 37 2084 21 98 Kam-24 5711 132 16 0.68 0.1080 1.3 0.9132 1.8 0.0613 1.2 0.74 659 17 661 16 651 51 102 Kam-25 3680 86 9 0.48 0.0959 1.1 0.7957 1.5 0.0602 1.1 0.72 594 14 591 12 609 46 97 Kam-26 7305 174 21 0.86 0.1011 1.2 0.8437 1.5 0.0605 0.9 0.82 621 14 621 14 622 37 100 Kam-27 9755 199 26 1.99 0.1018 1.2 0.8564 1.7 0.0610 1.2 0.70 628 16 625 14 639 52 98 Kam-28 8995 237 25 0.62 0.0945 1.2 0.7728 1.5 0.0593 0.9 0.79 581 13 582 13 578 40 101 Kam-29 62038 140 66 0.60 0.4052 1.0 7.639 1.5 0.1367 1.1 0.68 2190 27 2193 37 2186 38 100 Kam-30 4117 94 12 0.98 0.1018 1.2 0.8521 1.8 0.0607 1.3 0.69 626 17 625 15 629 55 99 Kam-31 7954 173 24 1.38 0.1041 0.8 0.8845 1.2 0.0616 0.9 0.65 643 11 638 9 661 39 97 Kam-32 3988 55 7 0.80 0.1098 0.7 0.9465 1.4 0.0625 1.2 0.52 676 14 671 10 692 53 97 Kam-33 5413 132 17 1.10 0.1030 0.8 0.8547 1.8 0.0602 1.6 0.44 627 17 632 9 610 69 104 Kam-34 4202 112 11 0.31 0.0956 0.9 0.7930 1.8 0.0602 1.6 0.49 593 16 589 10 609 68 97 Kam-35 4412 123 12 0.58 0.0918 0.8 0.7536 1.4 0.0595 1.1 0.60 570 12 566 9 586 47 97 Kam-36 10866 311 29 0.28 0.0928 0.8 0.7512 1.1 0.0587 0.8 0.67 569 10 572 8 556 36 103 Kam-37 5386 118 14 0.56 0.1091 0.8 0.9149 1.5 0.0608 1.2 0.57 660 14 667 11 634 51 105 Kam-38 3379 85 13 1.71 0.0999 0.8 0.8255 1.7 0.0599 1.5 0.49 611 16 614 10 600 64 102 Kam-39 3660 104 10 0.46 0.0866 0.9 0.7060 1.5 0.0591 1.2 0.59 542 13 535 9 572 52 94 Kam-40 3327 44 5 0.88 0.0924 1.2 0.7713 3.0 0.0606 2.8 0.39 581 27 570 13 623 120 91 Kam-41 4189 100 12 0.63 0.1070 0.8 0.9018 1.3 0.0611 1.1 0.60 653 13 655 10 644 46 102 Kam-42 3914 60 7 0.68 0.0971 0.8 0.8066 1.8 0.0602 1.6 0.46 601 16 597 9 612 68 98
Sample
Isotopic ratios§
TABLE 2. LASER ABLATION-ICP-MS U, Pb, AND Th DATA OF DETRITAL ZIRCON GRAINS FROM PALAEOZOIC QUARZITES OF THE ELBEZONE, SAXO-THURINGIAN ZONE, BOHEMIAN MASSIF
76 Linnemann et al.
Pb* (cps)
129 32 50 141 16 91 37 66 172 196 352 133 172 207 81 39 756
U (ppm)
†
14 15 6 14 2 12 4 9 30 24 44 52 19 23 8 15 77
Pb (ppm)
†
0.77 1.04 0.98 0.63 0.86 1.38 0.39 1.59 0.35 0.99 1.18 0.38 0.59 0.52 0.56 0.30 0.58
Th U
†
Pb U
0.0966 0.3675 0.0951 0.0889 0.1051 0.0992 0.0989 0.1080 0.1694 0.1036 0.0978 0.3615 0.0997 0.1015 0.0906 0.3627 0.0950
238
206
0.8 0.8 0.9 1.0 1.0 0.8 0.9 0.8 0.9 0.9 0.8 1.2 0.9 0.9 0.8 0.8 0.9
1σ (%)
Pb U
0.7979 6.424 0.7866 0.7158 0.8713 0.8177 0.8306 0.9196 1.727 0.8820 0.8109 6.208 0.8244 0.8525 0.7467 6.457 0.7857
235
207
1.4 1.3 1.6 1.6 2.8 1.8 1.9 1.6 1.1 1.2 1.0 1.3 1.6 1.3 2.5 1.3 1.0
1σ (%)
Pb Pb
0.0599 0.1268 0.0600 0.0584 0.0601 0.0598 0.0609 0.0618 0.0739 0.0617 0.0602 0.1245 0.0600 0.0609 0.0598 0.1291 0.0600
206
207
1.1 1.0 1.3 1.3 2.6 1.7 1.7 1.4 0.7 0.8 0.7 0.6 1.4 1.0 2.4 1.0 0.6
1σ (%)
0.60 0.63 0.55 0.59 0.34 0.44 0.47 0.46 0.79 0.73 0.75 0.88 0.54 0.68 0.31 0.63 0.82
Rho**
Pb U
596 2036 589 548 636 607 614 662 1019 642 603 2006 610 626 566 2040 589
235
207
13 24 14 14 27 17 18 16 15 11 10 23 15 12 22 23 9
±2σ (Ma)
Pb U
Ages
594 2018 586 549 644 610 608 661 1009 636 601 1989 613 623 559 1995 585
238
206
10 29 10 10 12 9 10 9 17 10 9 40 10 11 8 28 10
±2σ (Ma) Pb Pb
601 2054 603 545 608 597 636 666 1040 664 609 2022 602 636 595 2086 602
206
207
49 37 58 57 114 72 72 62 28 35 30 23 59 42 105 35 26
±2σ (Ma)
99 98 97 101 106 102 96 99 97 96 99 98 102 98 94 96 97
Conc (%)
Continued
Langerberg Quartzite (Location: microconglomerate, Lower Ordovician, Tremadoc, Langer Berg Formation, Schwarzburg antiform [northwestern part], Langer Berg near Willmersdorf, Easting: 64 3489, Northing: 56 09502) Lbq_1 103842 247 95 0.03 0.3873 1.3 7.467 1.8 0.1398 1.2 0.74 2169 32 2110 48 2225 41 95 Lbq_2 11930 148 15 0.51 0.1058 1.0 0.8833 4.1 0.0605 4.0 0.25 643 40 648 12 623 172 104 Lbq_3 10574 239 22 0.62 0.0958 1.0 0.7745 1.3 0.0587 0.8 0.78 582 11 590 11 554 35 106 Lbq_4 26758 107 34 0.57 0.2979 1.2 4.528 1.4 0.1102 0.7 0.85 1736 23 1681 34 1803 26 93 Lbq_5 52243 168 63 0.42 0.3476 1.0 5.628 1.1 0.1174 0.5 0.91 1920 20 1923 35 1917 17 100 Lbq_6 15342 360 32 0.49 0.0910 1.0 0.7369 1.1 0.0587 0.6 0.85 561 10 562 11 557 26 101 Lbq_7 217112 157 149 1.08 0.7058 1.2 28.37 1.3 0.2915 0.6 0.89 3432 26 3443 62 3426 19 101 Lbq_8 5306 110 11 0.45 0.1011 1.0 0.8419 1.9 0.0604 1.6 0.53 620 18 621 12 617 69 101 Lbq_9 24729 70 28 0.80 0.3720 1.1 6.275 1.2 0.1223 0.6 0.87 2015 22 2039 37 1991 22 102 Lbq_10 14014 331 28 0.14 0.0861 1.3 0.6943 1.5 0.0585 0.9 0.83 535 13 532 13 548 38 97 Lbq_11 17862 404 40 0.42 0.0927 1.2 0.7409 1.4 0.0580 0.7 0.88 563 12 571 13 529 29 108 Lbq_12 2644 62 6 0.42 0.0930 1.1 0.7542 1.8 0.0588 1.5 0.58 571 16 573 12 561 65 102 Lbq_13 3910 77 8 0.45 0.0922 1.0 0.7727 3.5 0.0608 3.4 0.29 581 31 568 11 632 145 90 Lbq_14 34784 864 70 0.31 0.0795 1.2 0.6387 1.3 0.0583 0.7 0.87 502 11 493 11 539 29 91 Lbq_15 10562 239 24 0.25 0.1003 1.1 0.8228 1.5 0.0595 1.0 0.74 610 14 616 13 585 43 105 Lbq_16 11153 247 25 0.43 0.0958 1.0 0.7707 1.3 0.0584 0.8 0.78 580 12 590 12 544 36 108 Lbq_17 5830 111 11 0.32 0.1045 1.1 0.9071 1.9 0.0630 1.6 0.56 655 18 640 13 707 66 91 Lbq_18 13751 300 30 0.66 0.1030 1.0 0.8443 1.3 0.0594 0.8 0.80 622 12 632 13 584 34 108 Lbq_19 119431 353 128 0.15 0.3588 1.2 5.811 1.3 0.1175 0.7 0.87 1948 23 1977 39 1918 24 103 Lbq_20 30878 76 35 0.74 0.3795 1.0 6.896 1.1 0.1318 0.5 0.90 2098 20 2074 35 2122 17 98 Lbq_21 57953 75 43 0.39 0.4966 1.0 13.65 1.1 0.1994 0.5 0.87 2726 21 2599 42 2821 18 92 Lbq_22 55794 145 62 0.55 0.3761 1.0 6.536 1.1 0.1260 0.5 0.90 2051 20 2058 36 2043 18 101
Kam-1 (continued) Kam-43 4659 Kam-44 9519 Kam-45 2073 Kam-46 7097 Kam-47 3630 Kam-48 3850 Kam-50 3271 Kam-51 2673 Kam-52 15153 Kam-53 8361 Kam-54 18610 Kam-55 37300 Kam-56 7154 Kam-57 7995 Kam-58 3800 Kam-59 6430 Kam-60 27906
Sample
207
Isotopic ratios§
TABLE 2. LASER ABLATION-ICP-MS U, Pb, AND Th DATA OF DETRITAL ZIRCON GRAINS FROM PALAEOZOIC QUARZITES OF THE ELBEZONE, SAXO-THURINGIAN ZONE, BOHEMIAN MASSIF (continued)
The continuum between Cadomian orogenesis and opening of the Rheic Ocean 77
0.44 0.59 0.66 0.28 0.42 0.34 0.08 0.61 0.32 0.89 0.46 0.52 0.32 0.64 0.40 0.72 0.28 0.54 0.61 0.47 0.50 0.33 0.43 0.02 1.14 0.46 0.48 0.59 0.62 0.79 0.66 0.40 0.28 0.51 0.82 1.18 0.83
Langerberg Quartzite (continued) Lbq_23 146269 329 148 Lbq_24 1268 31 3 Lbq_25 6566 118 12 Lbq_26 148581 182 109 Lbq_27 12729 265 29 Lbq_28 129573 259 117 Lbq_29 64458 91 47 Lbq_30 40378 964 96 Lbq_31 5515 62 9 Lbq_32 9340 184 22 Lbq_33 15446 339 33 Lbq_34 14116 262 27 Lbq_35 63888 617 65 Lbq_36 30832 97 39 Lbq_37 226574 167 127 Lbq_38 24408 570 58 Lbq_39 10780 120 14 Lbq_40 27007 529 58 Lbq_41 8050 194 19 Lbq_42 22501 272 28 Lbq_43 84932 236 92 Lbq_44 8245 26 11 Lbq_45 20607 136 12 Lbq_46 370489 488 227 Lbq_47 131963 173 118 Lbq_48 29502 72 30 Lbq_49 31150 701 64 Lbq_50 6911 142 15 Lbq_51 16378 359 37 Lbq_53 3977 84 10 Lbq_54 10025 181 18 Lbq_55 15696 330 32 Lbq_56 25267 577 52 Lbq_57 259062 419 215 Lbq_58 13896 152 29 Lbq_59 7482 156 20 Lbq_60 3088 64 7 0.4081 0.0915 0.1024 0.5386 0.1021 0.4159 0.4964 0.0882 0.1461 0.0989 0.0916 0.0946 0.0853 0.3407 0.6366 0.0886 0.0998 0.1012 0.0884 0.0955 0.3473 0.3971 0.0812 0.4461 0.5117 0.3665 0.0880 0.0964 0.0930 0.0965 0.0880 0.0917 0.0887 0.4447 0.1607 0.1001 0.0873
Pb 238 U
206
1.0 1.0 1.0 1.0 1.0 1.0 1.1 1.1 1.5 0.8 0.9 0.8 1.0 0.9 0.8 0.9 0.8 0.8 0.7 0.7 0.6 0.9 0.9 1.0 0.8 0.7 1.2 0.7 0.9 0.7 1.2 0.8 0.8 1.0 0.9 0.7 0.8
1σ (%)
7.611 0.7388 0.8639 14.06 0.8319 7.942 12.92 0.7244 1.409 0.8140 0.7523 0.7761 0.6956 5.749 22.40 0.7189 0.8294 0.8578 0.7173 0.7759 7.187 0.656 0.6557 12.01 13.09 6.930 0.7137 0.8044 0.7566 0.8032 0.7037 0.7641 0.7167 10.14 1.577 0.8345 0.7041
Pb 235 U
207
1.1 2.4 1.4 1.1 1.5 1.2 1.3 1.3 2.0 1.1 1.2 1.2 2.1 1.1 0.9 1.1 1.0 0.9 1.1 2.1 0.7 1.4 2.0 1.1 0.8 0.9 1.7 1.2 1.2 1.3 1.8 1.2 1.0 1.1 1.2 1.3 1.8
1σ (%)
0.1352 0.0585 0.0612 0.1894 0.0591 0.1385 0.1888 0.0595 0.0700 0.0597 0.0596 0.0595 0.0591 0.1224 0.2552 0.0588 0.0602 0.0615 0.0589 0.0589 0.1225 0.1313 0.0586 0.1952 0.1855 0.1371 0.0588 0.0605 0.0590 0.0604 0.0580 0.0605 0.0586 0.1654 0.0712 0.0605 0.0585
Pb 206 Pb
207
0.4 2.2 1.0 0.4 1.1 0.6 0.7 0.7 1.4 0.8 0.9 0.8 1.8 0.6 0.5 0.7 0.7 0.5 0.8 2.0 0.4 1.1 1.8 0.4 0.3 0.6 1.2 1.0 0.9 1.1 1.3 0.8 0.6 0.6 0.9 1.1 1.6
1σ (%)
0.92 0.40 0.72 0.94 0.68 0.87 0.86 0.86 0.73 0.75 0.70 0.70 0.48 0.83 0.87 0.79 0.76 0.83 0.63 0.33 0.87 0.65 0.45 0.95 0.92 0.79 0.73 0.58 0.71 0.55 0.68 0.71 0.79 0.85 0.68 0.56 0.42
Rho**
2186 562 632 2754 615 2224 2674 553 893 605 570 583 536 1939 3201 550 613 629 549 583 1956 2135 512 2605 2686 2103 547 599 572 599 541 576 549 2448 961 616 541
Pb 235 U
207
19 21 14 20 14 22 25 11 24 10 11 11 17 19 18 10 9 9 9 19 13 25 16 20 16 17 14 11 11 12 15 11 8 21 16 12 15
±2σ (Ma)
Ages
2206 565 628 2777 627 2242 2598 545 879 608 565 583 528 1890 3175 547 614 621 546 588 1921 2156 503 2378 2664 2013 544 593 574 594 544 565 548 2371 961 615 540
Pb 238 U
206
37 10 12 46 12 39 48 12 24 10 9 9 10 30 40 9 9 9 7 8 21 34 9 41 33 26 13 8 10 8 13 9 8 38 15 9 8
±2σ (Ma) Pb Pb
2167 550 647 2737 571 2208 2732 587 927 594 588 585 572 1991 3217 561 612 657 562 564 1993 2115 550 2787 2703 2191 561 623 566 616 529 620 553 2511 963 620 547
206
207
15 97 43 12 49 20 22 28 57 33 37 37 79 22 14 31 29 23 37 86 13 38 77 12 11 20 51 43 38 48 58 37 27 20 37 47 71
±2σ (Ma)
102 103 97 101 110 102 95 93 95 102 96 100 92 95 99 98 100 95 97 104 96 102 91 85 99 92 97 95 101 96 103 91 99 94 100 99 99
Conc (%)
Notes: Coordinates are UTM World Geodetic System 84. Conc.—concordance. *Within-run background-corrected mean 207Pb signal in counts per second. † U and Pb content and Th/U ratio were calculated relative to GJ-1 and are accurate to ~10%. § Corrected for background, mass bias, laser induced U-Pb fractionation and common Pb (if detectable; see text on analytical method) using the Stacey and Kramers (1975) model Pb composition. 207Pb/235U calculated using 207Pb/206Pb/(238U/206Pb × 1/137.88). Errors are propagated by quadratic addition of within-run errors (1 standard error) and the reproducibility of GJ-1 (1 standard deviation). **Rho is the error correlation defined as err206Pb/238U/err207Pb/235U. See text for details.
U† (ppm)
Th† U
Pb* (cps)
207
Pb† (ppm)
Sample
Isotopic ratios§
TABLE 2. LASER ABLATION-ICP-MS U, Pb, AND Th DATA OF DETRITAL ZIRCON GRAINS FROM PALAEOZOIC QUARZITES OF THE ELBEZONE, SAXO-THURINGIAN ZONE, BOHEMIAN MASSIF (continued)
78 Linnemann et al.
Pb* (cps)
U (ppm)
†
Pb (ppm)
†
Th U
†
Pb U
238
206
1σ (%) Pb U
235
207
1σ (%)
Isotopic ratios§ 206
207
Pb Pb
1σ (%)
Rho** Pb U
235
207
±2σ (Ma)
±2σ (Ma)
Ages Pb U
238
206 206
207
Pb Pb
±2σ (Ma)
Conc (%)
11174 15023 18762 15110 6612 11321 32390 21475 10974 6512 17517 17312 44335 6410 24741
291 407 505 317 178 270 870 602 262 181 523 459 1259 174 665
24 36 45 30 15 25 84 57 23 15 45 40 121 15 62
0.41 0.71 0.71 0.76 0.55 0.79 0.53 0.96 0.51 0.53 0.57 0.59 0.27 0.48 0.82
0.0798 0.0798 0.0800 0.0816 0.0798 0.0802 0.0813 0.0802 0.0804 0.0807 0.0802 0.0807 0.0788 0.0806 0.0800
0.6 0.5 0.7 0.8 0.8 0.6 0.7 0.5 0.8 0.8 0.8 0.8 0.8 0.9 0.9
0.6268 0.6335 0.6346 0.6496 0.6255 0.6254 0.6436 0.6340 0.6340 0.6312 0.6293 0.6404 0.6221 0.6385 0.6359
1.2 1.0 1.1 1.2 1.2 1.1 1.1 1.2 1.4 1.2 1.2 1.2 1.2 1.3 1.3
0.0570 0.0576 0.0576 0.0577 0.0569 0.0566 0.0574 0.0574 0.0572 0.0567 0.0569 0.0576 0.0572 0.0575 0.0577
1.1 0.9 0.8 0.9 1.0 0.9 0.8 1.1 1.2 0.9 0.9 0.8 1.0 0.9 0.9
0.51 0.54 0.63 0.68 0.63 0.59 0.66 0.45 0.58 0.65 0.63 0.73 0.63 0.71 0.73
494 498 499 508 493 493 505 499 499 497 496 503 491 501 500
10 8 8 9 10 8 9 9 11 10 9 9 10 10 10
495 495 496 506 495 497 504 497 499 500 497 500 489 500 496
6 5 7 8 8 6 7 5 8 8 7 8 8 9 9
491 515 513 520 487 475 506 505 499 481 488 514 500 510 518
47 37 37 37 42 39 36 47 51 41 41 35 43 40 39
101 96 97 97 102 105 100 98 100 104 102 97 98 98 96
5093 5303 6042 5438 6145 58778 10238 8023 7313 6387 4050 8081 5894 14059 2683 7614 4427 11356
127 125 126 152 265 79 290 333 241 160 134 297 167 342 71 210 119 318
11 11 12 13 21 46 22 17 20 15 9 15 15 30 6 17 10 23
0.64 0.65 0.68 0.76 0.59 0.68 0.14 0.77 1.24 0.28 0.64 0.42 0.73 0.64 0.70 0.43 0.66 0.41
0.0773 0.0790 0.0795 0.0746 0.0784 0.4927 0.0800 0.0412 0.0772 0.0934 0.0623 0.0491 0.0790 0.0791 0.0791 0.0772 0.0792 0.0707
1.8 1.4 1.5 1.7 1.6 1.4 1.7 3.2 2.1 1.6 1.6 2.5 1.4 2.0 1.9 1.3 1.5 1.5
0.6029 0.6245 0.6197 0.5845 0.6180 11.98 0.6294 0.3164 0.6086 0.7593 0.4888 0.3764 0.6230 0.6249 0.6264 0.5972 0.6242 0.5401
2.4 2.3 1.8 2.1 1.9 1.6 2.1 3.3 2.5 2.2 2.0 3.0 1.8 3.1 2.3 1.9 1.7 2.3
0.0566 0.0573 0.0566 0.0569 0.0572 0.1764 0.0570 0.0558 0.0572 0.0589 0.0569 0.0557 0.0572 0.0573 0.0574 0.0561 0.0572 0.0554
1.6 1.9 1.0 1.3 1.1 0.8 1.3 1.0 1.3 1.5 1.3 1.6 1.1 2.4 1.3 1.4 0.9 1.8
0.74 0.61 0.83 0.80 0.81 0.85 0.79 0.96 0.85 0.73 0.76 0.84 0.78 0.63 0.82 0.67 0.84 0.64
479 493 490 467 489 2603 496 279 483 574 404 324 492 493 494 475 492 438
18 18 14 16 15 30 17 16 19 19 14 17 14 24 18 15 14 16
480 490 493 464 486 2583 496 260 479 576 389 309 490 491 491 479 491 440
16 14 14 15 15 58 16 16 20 18 12 15 13 19 18 12 14 12
475 503 474 486 499 2619 493 443 498 565 489 439 500 503 507 458 498 429
70 82 44 56 50 28 58 43 59 65 58 73 49 106 59 64 41 79
101 98 104 95 98 99 101 59 96 102 80 70 98 98 97 105 99 103
Notes: Coordinates are UTM World Geodetic System 84. Conc.—concordance. *Within-run background-corrected mean 207Pb signal in counts per second. † U and Pb content and Th/U ratio were calculated relative to GJ-1 and are accurate to ~10%. § Corrected for background, mass bias, laser induced U-Pb fractionation and common Pb (if detectable; see text on analytical method) using the Stacey and Kramers (1975) model Pb composition. 207Pb/235U calculated using 207Pb/206Pb/(238U/206Pb × 1/137.88). Errors are propagated by quadratic addition of within-run errors (1 standard error) and the reproducibility of GJ-1 (1 standard deviation). **Rho is the error correlation defined as err206Pb/238U/err207Pb/235U. See text for details.
kArc-16 kArc-17 kArc-18 kArc-19 kArc-20 kArc-21 kArc-22 kArc-23 kArc-24 kArc-25 kArc-26 kArc-27 kArc-28 kArc-29 kArc-30 kArc-31 kArc-32 kArc-33
kArc1 (Location: pebble-bearing rhyolithic tuffite (traditional term in the older German references: “Konglomeratische Arkose”), Lower Ordovician, Tremadocian, Goldisthal Formation, Schwarzburg antiform [southeastern part], valley of the Blambach near Sitzendorf, Easting: 65 2447, Northing: 56 11732)
Ves1_1 Ves1_2 Ves1_3 Ves1_4 Ves1_5 Ves1_6 Ves1_7 Ves1_8 Ves1_9 Ves1_10 Ves1_11 Ves1_12 Ves1_13 Ves1_14 Ves1_15
Ves1 (Location: rhyolitic ignimbrite, Upper Cambrian, Rollkopf Formation, Vesser complex, old quarry close to the ski-jump at Schmiedefeld, Easting: 62 8631, Northing: 56 06911)
Sample
207
TABLE 3. LASER ABLATION-ICP-MS U, PB, AND TH DATA OF ZIRCON FROM PYROCLASTIC SEDIMENTS OF THE SAXO-THURINGIAN ZONE, BOHEMIAN MASSIF
The continuum between Cadomian orogenesis and opening of the Rheic Ocean 79
80
Linnemann et al.
was used to enable sequential sampling of heterogeneous grains (e.g., growth zones) during time-resolved data acquisition (see also Janoušek et al., 2006). Each analysis consisted of ~20 s background acquisition followed by 35 s data acquisition, using a laser spot size of 30 and 40 μm, respectively. A common-Pb correction based on the interference- and background-corrected 204Pb signal and a model Pb composition (Stacey and Kramers, 1975) was carried out if necessary. The necessity of the correction is based on whether the corrected 207Pb/206Pb lies outside of the internal errors of the measured ratios. Discordant analyses were generally interpreted with care. Raw data were corrected for background signal, common Pb, laser-induced elemental fractionation, instrumental mass discrimination, and time-dependant elemental fractionation of Pb/Th and Pb/U using an Excel® spreadsheet program. Reported uncertainties were propagated by quadratic addition of the external reproducibility (standard deviation) obtained from the standard zircon GJ-1 (n = 18; ~0.6% and 0.5–1.0% for the 207 Pb/206Pb and 206Pb/238U, respectively) during individual analytical sessions and the within-run precision of each analyses (standard error). Concordia diagrams (2σ error ellipses) and concordia ages (95% confidence level) were produced using Isoplot/Ex 2.49 (Ludwig, 2001) and frequency and relative probability plots using AgeDisplay (Sircombe, 2004). The 207Pb/206Pb age was taken for interpretation for all zircons >1.0 Ga, and the 206Pb/238U ages was used for younger grains. For further details on analytical protocol and data processing see Gerdes and Zeh (2006).
data-point error ellipses are 2σ 206
0.6
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Pb
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n = 61
3000
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1.2
RESULTS 1.0
The results of LA-ICP-MS U-Pb zircon dating are listed in Tables 1–3 and shown on the concordia diagrams in Figures 7, 9, and 11 (see below). Binned frequency and probability density distribution plots are shown in Figures 8 and 10. For the latter two plots only those analyses less than 10% discordant were used. In this study, all systems, erathems, and eonothems are used in accordance with the stratigraphic table of Gradstein et al., (2004). Percentages of zircon ages for each sample are shown in Table 4. For zircons older than 1.0 Ga, the 207Pb/206Pb age is mentioned in the text. Younger zircon ages refer to the 206Pb/238U age. From sample Pur-1, seventy-four detrital zircon grains were analyzed, sixty-seven of which yielded concordant ages (Figs. 7 and 8). The age of the youngest concordant grain is 558 ± 16 Ma, and that of the oldest is 3465 ± 20 Ma (Table 1). Archean ages make up ~21% of the population. These ages fall into two groups at 3.5–3.3 Ga (six grains) and at 2.9–2.6 Ga (nine grains). About 33% of the grains record Paleoproterozoic ages between 1.8 and 2.3 Ga and nearly 45% record Neoproterozoic ages. The later group show pronounced peaks at ca. 640, ca. 620–590, and ca. 570 Ma and less pronounced peaks at ca. 710, ca. 790, and ca. 1000–950 Ma (Fig. 8, Table 4). All sixty-one analyses from sample Roth-1 are 90–110% concordant (Figs. 7 and 8). The ages range from 552 ± 11 Ma to 3053 ± 11 Ma. Only 5% of the analyzed grains are Archean in age, ~20% Paleoproterozoic, and ~75% Neoproterozoic.
0
4 206
Pb
238
0.6
8
12
16
Pur-1
20
n = 74
24
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U 2800
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0
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16
Pb/
235
20
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Figure 7. Concordia plots of LA-ICP-MS U-Pb analyses of detrital zircon grains from Late Neoproterozoic (Ediacaran) sedimentary rocks of the Saxo-Thuringian zone: Purpurberg quartzite from the Weesenstein Group (Pur-1), microconglomerate from the Lausitz Group (Wett-1), and graywacke from the Rothstein Formation (Roth-1). Error ellipses are 2σ. Insets show enlargement of the younger ages. n—number of analyses. For sample details see Table 1.
14
10
81
n=15/15 90–110% conc.
3400
3200
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900
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n=46/46 90–110% conc.
500
Frequency
12
Relative Probability
The continuum between Cadomian orogenesis and opening of the Rheic Ocean
6
Roth-1 (Neoproterozoic)
4
n=61/61, 90-110% conc.
2
16 14
10
N=20/22 90–110% conc.
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n=39/39 90–110% conc.
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Frequency
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Relative Probability
0
Wett-1 (Neoproterozoic)
4
n=59/61, 90-110% conc. 2
10 9 8
6
3400
3200
3000
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1000
900
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3
700
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5
n=36/42 90–110% conc.
n=31/32 90–110% conc.
500
Frequency
7
Relative Probability
0
Pur-1 (Neoproterozoic) n=67/74, 90-110% conc.
2 1 0 3400 3300 3200 3100 3000 2900 2800 2700 2600 2500 2400 2300 2200 2100 2000 1900 1800 1700 1600 1500 1400 1300 1200 1100 1000 900 800 700 600 500
Age (Ma) Figure 8. Binned frequency and probability density distribution plots of detrital zircon grains from Late Neoproterozoic (Ediacaran) sandstones of the Saxo-Thuringian zone. conc.—concordance; n—number of analyses with <10% discordance/total number of analyzed grains.
8% 0 0 0 3% 13 % 5% 10 % 0 9% 33 % 20 % 24 % 15 % 23 % 3% 1% 7% 2% 4% 3% 8% 3% 0 0 Note: The total range of measured U-Pb ages are shown in parentheses.
40 % 66 % 48 % 83 % 47 % 0 0 8% 0 5% 0 0 0 0 9% Neoproterozoic Neoproterozoic Neoproterozoic Lower Cambrian Lower Ordovician Pur-1 Roth-1 Wett-1 Kam-1 Lbq-1
67 61 59 58 57
(ca. 3220– 3465 Ma) (ca. 2510– 3180 Ma) (ca. 1750– 2480 Ma) (ca. 900– 1050 Ma) (ca. 700– 790 Ma) (ca. 540– 545 Ma) (ca. 490– 535 Ma)
(ca. 550– 690 Ma)
Liberian Eburnean Grenvillian Rodinia break-up Cadomian back-arc Cadomian retroarc
Sample Sample
Rifting
NeoMesoproterozoic Neo- to transition interval Paleoproterozoic Mesoarchean NeoproterozoicCambrian transition Late Midinterval Neoproterozoic Neoproterozoic Sediment age
TABLE 4. POPULATIONS AND PERCENTAGES OF ALL ZIRCON CONCORDANT BETWEEN 90% AND 110%
Total number of grains
Leonian
Linnemann et al. Paleoarchean
82
Archean ages (three grains) range from 3.05 to 2.6 Ga, Paleoproterozoic ages from 2.17 to 1.78 Ga, and the Neoproterozoic ages fall mainly in the interval 650–550 Ma. There are clear clusters at ca. 612 (nine grains) and ca. 642 Ma (six grains), whereas the age range 600–550 Ma is defined by a broad peak (Fig. 8). In addition, minor peaks occur at 710, ca. 740, ca. 790, and ca. 960 Ma, similar to those of Pur-1. From sample Wett-1, sixty-one zircon grains were analyzed, of which fifty-nine are 90–110% concordant (Figs. 7 and 8). The ages vary from 3181 to 542 ± 27 Ma, of which ~10% are Archean, 24% Paleoproterozoic, and ~60% Neoproterozoic. Archean ages range from 3.2 to 2.6 Ga with a cluster at ca. 2.8– 2.6 Ga (five grains). Paleoproterozoic ages (fifteen grains) define several peaks in the range 2.17–1.76 Ga, whereas Neoproterozoic ages fall mostly in the interval 650–540 Ma, with five relatively well-defined clusters at ca. 642 (six grains), 625–610 (five grains), ca. 590 (ten grains), ca. 572 (five grains), and ca. 543 Ma (five grains). The latter subpopulation yields a concordia age of 543 ± 4 Ma (see Fig. 11A), which straddles the PrecambrianCambrian boundary (542 ± 1 Ma; Bowring et al., 2003). From the Cambrian sandstone Kam-1 we analyzed fifty-nine zircons, only one of which is more than 10% discordant (Figs. 9 and 10). No Archean and only nine Paleoproterozoic grains (15%) were identified, although the age of the oldest grain (2478 ± 16 Ma), which is close to the Archean boundary, must be considered a minimum age. Paleoproterozoic ages fall mostly in the interval 2.1–2.0 Ga, with one grain yielding a 206Pb/238U age of 1000 ± 18 Ma. Neoproterozoic ages, which make up more than 80% of the population, fall predominantly in the interval 690– 550 Ma. The latter population shows clear clusters at ca. 570, ca. 590, ca. 610, ca. 623, ca. 643, and ca. 665 Ma, each of which is represented by six to nine grains. Two grains yielded ages of ca. 547 Ma, and one grain yielded a 206Pb/238U age of ca. 535 ± 9 Ma. The latter analysis is ~6% discordant but is considered to provide a maximum depositional age constraint. From the Lower Ordovician microconglomerate Lbq-1 we analyzed fifty-nine zircons; fifty-seven analyses yielded 90–110% concordant ages, of which 12% are Archean (seven grains), 23% Paleoproterozoic (thirteen grains), and 60% Neoproterozoic to Cambrian. The Archean grains yielded ages of ca. 2511, ca. 2820–2700 (four grains), ca. 3217, and ca. 3246 Ma. The Paleoproterozoic ages define six peaks in the interval 2200– 1800 Ma (Fig. 10) and Neoproterozoic ages (thirty-three grains) fall predominantly in the interval 650–540 Ma, with a broader peak at ca. 636 Ma and clear clusters at ca. 615, ca. 590, ca. 570, and ca. 546 Ma, each defined by five to seven analyzed grains. The latter subpopulation of seven grains yielded a concordia age of 546 ± 4 Ma, which straddles the Precambrian-Cambrian boundary like that of Wett-2. In addition, two grains yielded ages close to the Neo-Mesoproterozoic boundary (927 ± 57 and 961 ± 16 Ma), and two others gave a Cambrian concordia age of 530 ± 8 Ma. The two youngest grains (503 ± 8 and 493 ± 12 Ma) straddle the Cambrian-Ordovician boundary close to the deposition age of the sample. The Th/U ratio of most of the measured
data-point error ellipses are 2σorogenesis and opening of the Rheic Ocean The continuum between Cadomian
206
Pb
238
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n=20/21 90–110% conc.
n=37/38 90–110% conc.
400
Frequency
7
Relative Probability
Figure 9. Concordia plots of LA-ICP-MS U-Pb analyses of detrital zircon grains from Lower Cambrian sandstone (Kam-1) and Early Ordovician microconglomerate (Lbq-1) of the Saxo-Thuringian zone. Error ellipses are 2σ. Insets show enlargement of the younger ages. n—number of analyses. For sample details see Table 2.
Lbq-1 (Lower Ordovician)
3
n=57/59, 90-110% conc.
2 1
14
10
n=9/9 90–110% conc.
3400
3200
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n=49/49 90–110% conc.
400
Frequency
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Relative Probability
0
6 4
Kam-1 (Lower Cambrian) n=58/59, 90-110% conc.
2 0 3400 3300 3200 3100 3000 2900 2800 2700 2600 2500 2400 2300 2200 2100 2000 1900 1800 1700 1600 1500 1400 1300 1200 1100 1000 900 800 700 600 500 400
Age (Ma)
Figure 10. Binned frequency and probability density distribution plots of detrital zircon grains from sample Kam-1 (Lower Cambrian) and Lbq-1 (Early Ordovician). conc—concordance; n— number of analyses with <10% discordance/total number of analyzed grains.
84
Linnemann et al.
zircons vary between 0.1 and 1.0 (Tables 1–3), a range that is typical for zircon crystallized from magmas of intermediate to felsic composition (e.g., Hoskin and Schaltegger, 2003). From the Vesser complex ignimbrite Ves-1 we analyzed fifteen zircon grains, which yielded a concordia age of 497 ± 2 Ma (Fig. 11B). Thus, the upper part of the Rollkopf Formation was most likely deposited during the uppermost Cambrian. From the pebble-bearing rhyolitic tuffite KArc-1, eighteen needle-shaped zircon grains were analyzed (Table 3). Nine of these yielded a concordia age of 486 ± 4 Ma, which is identical to the upper-intercept age defined by the four discordant analyses (Fig. 11C). The concordia age is interpreted to provide maximum age constraints for tuffite deposition. In addition, one zircon yielded a concordant Archean age of 2.62 Ga (Table 3). AGE OF SEDIMENTARY DEPOSTION In each sample, the youngest concordant zircon age provides a maximum constraint for the deposition of the sampled unit. In the case of the Purpurberg quartzite and the Rothstein Formation,
the youngest ages are 558 ± 16 Ma and 552 ± 11 Ma, respectively. However, the uncertainty in the degree of concordance of Neoproterozoic-Paleozoic grains dated by the LA-ICP-MS method is relatively large, and results obtained from just a single analysis have to be interpreted with care. For example, a typical uncertainty of 2–3% (2σ) in 207Pb/206Pb for a Late Neoproterozoic grain (e.g., 560 Ma) relates to an absolute error on the 207Pb/206Pb age of 44–65 Ma. Thus, the youngest grains in both samples can be grouped in the ca. 570-Ma age population. Seven grains that define this population in Pur-1 and Roth-1 define concordia ages of 570 ± 4 Ma (mean squared weighted deviated [MSWDC+E] = 0.59) and 566 ± 4 Ma (MSWDC+E = 0.95), respectively. These ages are in agreement with a SHRIMP age of 566 ± 10 Ma for a tuff from the Rothstein Formation (Buschmann et al., 2001). It is therefore likely that the ca. 570-Ma grains in Pur-1 and Roth-1 originated from tuff horizons and that their crystallization ages closely date the depositional ages of the Purpurberg quartzite, Weesenstein Group, and Rothstein Formation, respectively. In other words, the deposition of Weesenstein Group and the Rothstein Formation took place shortly after ca. 570 and ca. 566 Ma,
data-point error ellipses are 2 σ
data-point error ellipses are 2 σ 0.092
206
Pb 238 U
Wett-1
0.090
youngest zircon grains 560 only
0.085
Pb
238
Ves-1
U
0.083
550 0.088
206
510 500
0.081
540
490
0.079
0.086
543 ±4 Ma
530 0.68
MSWDC+E = 0.30 Probability C+E = 0.97 0.70 207
Pb/
0.72 235
0.74
0.077
520
497 ±2 Ma
480
MSWDC+E = 0.81 Probability C+E = 0.77
A
0.76 0.60
U
0.62
0.64
0.66
B
0.68
data-point error ellipses are 2 σ
Figure 11. Concordia plots showing LA-ICP-MS U-Pb ages important for the modeling of the Cadomian orogeny and the opening of the Rheic Ocean in the Saxo-Thuringian zone of the Bohemian Massif. Error ellipses are 2σ. For sample details see Tables 1 and 3. (A) Concordia age of 543 ± 4 Ma calculated from the ages of the five youngest detrital zircon grains of microconglomerate sample Wett-1, placing the maximum age of deposition of the Lausitz Group close to the Precambrian-Cambrian boundary (Lausitz antiform, Saxo-Thuringian zone). (B) Concordia age of 497 ± 2 Ma calculated from the ages of fifteen magmatic zircon grains from rhyolithic ignimbrite sample Ves-1 from the Vesser complex. (C) Concordia age of 486 ± 4 Ma obtained from the ages of nine magmatic zircon grains from pebble-bearing felsic tuffite sample Karc-1 from the base of the Ordovician in the southeastern part of the Schwarzburg antiform. The upper intercept age of 481 ± 15 Ma, defined by four discordant grains, overlaps with the concordia age. MSWDC+E—mean squared weighted deviates of concordance and equivalence.
520
206
0.08
Pb 238 U
KArc-1 440
0.07
400 0.06
360 320
0.05
280 0.04
481 ±15 Ma (upper intercept age) MSWD = 0.71
to zero 0.35
486 ±4 Ma MSWDC+E = 0.81
C 0.45
0.55
0.65
The continuum between Cadomian orogenesis and opening of the Rheic Ocean respectively. A Neoproterozoic age for both units is also indicated by the intrusion of the Dohna granodiorite into the Weesenstein Group at 537 ± 7 Ma (Linnemann et al., 2000; Fig. 3) and the overlaying of the Rothstein Formation by the Zwethau Formation during the Atdabanian (Lower Cambrian) at ca. 534 Ma (Buschmann et al., 2006). The youngest zircon subpopulation in sample Wett-1 yielded an age (543 ± 4 Ma) close to the Neoproterozoic-Cambrian boundary. This age (Fig. 11A) is interpreted to provide a maximum depositional age for the Lausitz Group. However, shortly after its deposition, the Lausitz Group was deformed and subsequently intruded by voluminous granitoids at 539 ± 6 Ma (Linnemann et al., 2000). Given the uncertainties in the age determinations, deposition of the Lausitz Group can be constrained to a narrow time interval of ~10 m.y. (between 547 and 533 Ma) around the Precambrian-Cambrian boundary. The control for the age of deposition of the Zwethau Formation (Kam-1) comes from paleontological data. Based on trilobites and archaeocyatha, the formation is assigned to the Atdabanian, that is, the Lower Cambrian stage starting at ca. 534 Ma (Elicki, 1997, and references therein). The depositional age of the Langer Berg Formation (Lower Ordovician, Lbq-1) is controlled by trace fossils and regional lithostratigraphic correlation, with well-dated sections in the southeastern part of the Schwarzburg antiform (Linnemann, 1996). In both samples the youngest zircon ages (Kam-1, ca. 535 Ma; Lbq-1, two grains, 498 ± 7 Ma) are very close to the biostratigraphic-controlled deposition age. The age of 497 ± 2 Ma (Fig. 11B) for Ves-1 zircons is interpreted to date ignimbrite deposition within the upper part of the Rollkopf Formation in the Vesser complex. This part of the Rollkopf Formation is therefore Upper Cambrian, according to Gradstein et al. (2004), who place the upper boundary of the Middle Cambrian at 501 ± 2 Ma. In contrast, the lower part of the Rollkopf Formation is assigned to the Middle Cambrian, based on conventional U-Pb dating of zircon from a dacitic pyroclastite (508 ± 2 Ma) and from a gabbro (502 ± 2 Ma) that intrudes the units (Kemnitz et al., 2002; see also Fig. 6). At present no biostratigraphic or geochronological data are available for the Neuwerk Formation in the upper part of the Vesser complex. The overlying Hundsrück Group contains felsic volcanic rocks and highly mature quartzites that are interpreted to be Lower Ordovician, based on lithostratigraphic correlation (Bankwitz et al., 1992; Linnemann, 2003a; Fig. 6). The age of 486 ± 4 Ma for seventeen zircon grains from sample KArc-1 is interpreted to closely date the deposition of the lithostratigraphic unit enclosing the tuffite. The sample was taken from the level at which the boundary between the Cadomian basement and Lower Paleozoic overstep sequence is suspected. This transition interval is free of key fossils. In this section of the southeastern part of the Schwarzburg antiform no angular unconformity between Cadomian basement and the Lower Paleozoic sediments is observed. The age of the KArc-1 zircons suggests that the Cadomian basement was directly overlain by highly mature sediments of lowermost Tremadocian age
85
(lowermost Ordovician; Fig. 6) and, thus, the entire Cambrian is missing. PROVENANCE OF SEDIMENTS The analyzed samples were selected to be representative of different successions of distinct ages. However, some care must be taken in the interpretation of the age spectra because a certain degree of sample bias cannot be excluded. Nevertheless, the U-Pb age spectra of the detrital zircons from samples Pur-1, Wett-1, Roth-1, Kam-1, and Lbq-1 show striking similarities (Figs. 8 and 10, Table 4), indicating that they display a common characteristic of the source area. Hence, variations among the samples are interpreted as indicating variations in the source area. All five samples predominantly contain Late Neoproterozoic (690–550 Ma; 40– 83%) and Paleoproterozoic (2.2–1.8 Ga; 15–33%) grains with a smaller fraction of Neo- to Mesoarchean constituents (5–13%; Table 4). In addition, all samples contain a small fraction of ca. 1000- to 900-Ma (1–7%) grains. Only the three Neoproterozoic sediments contain Mid-Neoproterozoic (790–700 Ma; 3–8%) zircons, and Paleoarchean (3–8%) components are present only in Pur-1 and Lbq-1. A common feature of all samples is an “age gap” between 1.75 and 1.0 Ga, which is typical of a Cadomian and/or west African provenance and is diagnostic in distinguishing it from East Avalonia and Baltica (e.g., Nance and Murphy, 1994; Friedl et al., 2000). This age gap is in agreement with the characteristic clusters of Paleoproterozoic ages in the interval 2.17–1.78 Ga. Such ages are typical of the western part of the Gondwana supercontinent, which was affected by abundant magmatic intrusions (ca. 2.2–1.8 Ga) during the Eburnean orogeny (West African craton). Furthermore, Neo- to Mesoarchean zircon ages and, in the case of Pur-1 and Lbq-1, additional Paleoarchean grains, point to recycling of magmatic rocks formed during the Liberian and the Leonian orgenies, respectively. Both orogenies affected the West African craton during the Archean. The overall dominance of Late Neoproterozoic zircon ages with apparent clusters at ca. 570, ca. 590, ca. 620–610, ca. 640, and ca. 660 Ma point to the fact that the most important source of sedimentary detritus (>55%) was the active magmatic arc of the Cadomian belt (ca. 700–550 Ma; e.g., Linnemann et al., 2000, 2004; Nance et al., 2002; Murphy et al., 2004). Taking all 165 Neoproterozoic zircon ages (700–550 Ma) into consideration, the main magmatic activity of this Cadomian arc took place at ca. 615, ca. 590, and ca. 570 Ma (28, 27, and 21%, respectively). Orthogneisses of Mid-Neoproterozoic protolith age (ca. 755–700 Ma) have been described from the Armorican Massif (Samson et al., 2003) and the Anti-Atlas orogen in Morrocco (D’Lemos et al., 2006). Based on Hf and Nd isotopes these have been interpreted as remnants of an earlier, independent arc system. Such ages are relative rare in the Avalonian-Cadomian belt. Our results suggest that such sources were present only during the Neoproterozoic and were absent during younger sedimentation. All samples contain a small fraction of grains with ages close to the Neo- to Mesoproterozoic boundary. Until recently,
86
Linnemann et al.
such “Grenville” ages were interpreted as being derived either from the Amazonian craton (e.g., Friedl et al., 2000; Hegner and Kröner, 2000; Fernández-Suárez et al., 2002) or from the Grenville orogenic belt (e.g., Nance and Murphy, 1994). Grenville ages typically fall in the range 1.25–1.0 Ga (e.g., Keppie et al., 1998), and an Amazonian provenance is characterized by various Mesoproterozoic age peaks in the 1.8- to 1.0-Ga interval (e.g., Nance and Murphy, 1994; Friedl et al., 2000). These ca. 1.05- to 0.9-Ga zircon ages are relatively rare but have been described from various localities within Cadomia (Gebauer et al., 1989; Fernández-Suárez et al., 2002; Gehmlich, 2003; Friedl et al., 2004). They therefore represent a typical Cadomian detrital zircon subpopulation and most likely also have a west African origin. The presented zircon ages strongly suggest a provenance
Peri-Gondwana Urals
ria
Carolina
Ibe
qui a Yucatan tis
Cor 0.9-1.2 1.3-2.2 2.5-2.85 3.0-3.2 Ga *4
ca. 570 Ma
SXZ TBU
MIA
CADO
0.54-0.7 1.8-2.2 2.7-2.9 3.1-3.4 Ga *1
"P ro
PLATE-TECTONIC MODEL The data set presented above provides several cornerstones that must be taken into account in the reconstruction of the Neoproterozoic to Cambro-Ordovician evolution of the SaxoThuringian zone and adjoining crustal units of the Bohemian Massif. First, all investigated units contain considerable quantities of Archean, Paleoproterozoic, and Neoproterozoic zircons, which provide evidence of intense crustal recycling (Linnemann et al., 2000, 2004; Drost et al., 2004; this study). The detrital age spectra point to a west African provenance and exclude an Ama-
Margin to -A
Turkish plate Aegean Dobrogea lp s" 0.55-0.65 0.9-1.1 1.65-1.85 2.45-2.7 Ga *2
Florida 0.54-0.7 1.0-1.35 1.45-1.75 2.5-3.1 Ga *3
uth So
W-A v
alo
nia
Av a
lo
E
va -A
AM MC ia F n lo
Oa xa
ni
an
-
C
ian m o ad
Active
of the investigated Neoproterozoic rocks from the margin of the West African craton (Fig. 12).
Pe ri-G on dw ana
Ea
st
Go nd wa na
po le
t
es W
an w nd o G
a
Neoproterozoic mobile belts of peri-Gondwana (Cadomian and related events) Neoproterozoic mobile belts of Gondwana (Pan-African and related events) 1.1 - 1.3 Ga Megashear event in Amazonia
Mesoproterozoic mobile belts (Grenville and related events) Cratons (Archean-Paleoproterozoic)
Figure 12. Paleogeography of the Cadomian-Avalonian active margin and related major peri-Gondwanan terranes at ca. 570 Ma. AM—Armorican Massif; FMC—French Massif Central; SXZ—Saxo-Thuringian zone (part of the Bohemian Massif); TBU—Teplá-Barrandian unit (part of the Bohemian Massif). Numbers in circles: zircon ages from the cratons in Ga. *1—from the compilation of Nance and Murphy (1994 and references therein); *2—from Avigad et al. (2003); *3—from Schneider Santos et al. (2000); *4—from the compilation of Zeh et al. (2001). Modified after Nance and Murphy (1994, 1996), Linnemann et al. (2000, 2004), Murphy et al. (2000), Linnemann and Romer (2002), Nance et al. (2002); paleogeography of the Gondwanan continental plates after Unrug (1996).
The continuum between Cadomian orogenesis and opening of the Rheic Ocean zonian and/or Baltic source (Drost et al., 2004; Linnemann et al. 2004; this study). Second, the zircon age spectra indicate that some Late Neoproterozoic sediments were deposited at about the same time (ca. 570 Ma), during, and after intense magmatic activity in the adjacent arcs. Third, most sedimentary basins were deformed between ca. 570 and 540 Ma and intruded by voluminous granitoid plutons at ca. 540 Ma. The absence of Cadomian high-grade or high-pressure regional metamorphic rocks like granulite and eclogite suggests that this structural-magmatic event took place without major thickening of the crust and without subduction of continental crust (Linnemann et al., 2000). Electron microprobe dating of metamorphic monazite grains from the Teplá-Barrandian unit yielded Th-U-Pb model ages of 540 ± 16, 542 ± 13, and 551 ± 19 Ma, which Zulauf et al. (1999) related to low-pressure/high-temperature metamorphism. Fourth,
MOR-basalts, andesites, and hydrothermal black cherts
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the generation of Lower to Middle Cambrian rift-related sediments, Upper Cambrian (ca. 500 Ma) MOR-related mafic rocks (Bankwitz et al., 1992; Kemnitz et al., 2002; this study), and thick Lower Ordovician successions with high subsidence rates indicate the formation of a rift-drift succession in Cambro-Ordovician time (Linnemann and Romer, 2002; Kroner et al., 2003; Linnemann et al., 2004). Cadomian Back-Arc Basin Evolution The oldest rocks of the Saxo-Thuringian zone are sediments deposited at ca. 570–565 Ma. Rock units belonging to this age interval make up the Weesenstein and Clanzschwitz groups as well as the Altenfeld and Rothstein Formations (Figs. 3 and 13). The Altenfeld Formation of the Schwarzburg antiform comprises
Closed strike-slip Clanzschwitz Group basin: redeposition Weesenstein of sediments and Group arc-related igneous rocks
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Intra-arc basin
Transitional crust: early Cadomian arc (c. 600-650 Ma)
Continent (crust of the West African craton)
Cadomian back-arc basin c. 590-560 Ma
Figure 13. Model for the plate-tectonic development of the Cadomian back-arc basin at ca. 590 and 560 Ma, based on data derived from the Saxo-Thuringian zone (Bohemian Massif). Back-arc basin consists of a continentward passive margin, represented by the Weesenstein and Clanzschwitz groups, and an arcward margin, characterized by more strongly stretched continental crust and the accumulation of predominantly arc-derived debris. The back-arc is documented by MOR-related rocks and hydrothermal black cherts recorded in the Altenfeld and the Rothstein formations. Inset: Sketch of analogous plate-tectonic configuration represented by the opening of the Japan Sea in the western Pacific region during the Early Miocene (after Jolivet et al., 1992). The back-arc basin of the Japan Sea is largely floored by stretched continental crust.
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similar facies to those of the Rothstein Formation (black cherts, mafic igneous rocks), suggesting a similar age and plate-tectonic setting of deposition (Fig. 13). In the same way the Clanzschwitz Group can be correlated with the Weesenstein Group (Pur-1). We assume that all of these units were deposited in a backarc basin, which predominantly consisted of thinned continental crust and was flanked by a magmatic arc to the “north” and by a cratonic source to the “south” (Linnemann et al., 2000; Buschmann et al., 2001; Figs. 3 and 13). Back-arc spreading at ca. 570 Ma is best documented by the Rothstein Formation (Buschmann, 1995). Field data and geochemical information suggest that the 566 ± 10-Ma-old Rothstein Formation comprises a low-grade metamorphic suite of intrusive and effusive enriched mid-ocean ridge basalts (E-MORB), andesites, calc-alkaline metabasalts, and subordinate alkaline metabasalts (Buschmann, 1995; Buschmann et al., 2001). The submarine effusive character of these rocks is indicated by pillow structures that may have formed during seafloor spreading. The submarine character is additionally supported by black cherts, which are assumed to be the product of hydrothermal activity at a spreading center that caused alteration of the submarine volcanic and sedimentary rocks (Fig. 4A). According to Buschmann (1995), deposition of the Rothstein Formation was accompanied by strike-slip faulting that produced submarine pull-apart basins and led to the re-sedimentation of older unconsolidated sediments. A similar age and tectonic regime is assumed for the Altenfeld Formation. In contrast to the Rothstein and Altenfeld formations, the Weesenstein and the Clanzschwitz groups were most likely deposited at the passive margin of the back-arc basin, the existence of which is indicated by (1) Nd-model ages for the sediments in the range ca. 2.1–1.5 Ga (Linnemann and Romer, 2002), (2) abundant Paleoproterozoic detrital zircon ages (Linnemann et al., 2004; this study), and (3) the presence of highly mature sediments like the Purpurberg quartzite (Linnemann, 1991; Fig. 13). The existence of a Cadomian magmatic arc can be deduced from the geochemical signatures of the Late Neoproterozoic sediments (Buschmann, 1995; Linnemann and Romer, 2002; Drost et al., 2004). These sediments have a felsic provenance pointing to a relatively mature continental arc with a relatively thick root zone. The main phase of arc magmatism occurred between ca. 560 and 600 Ma (Linnemann et al., 2004; this study). Comparable ages and arc-related igneous rocks are also described from the Avalonian part of the Bohemian Massif (Finger et al., 2000; Friedl et al., 2004; Fig. 2). Relatively small remnants of the Cadomian arc are known from the Armorican Massif and Iberia. Also in the Bohemian Massif arc remnants sensu stricto are scarce (e.g., Krˇíbek et al., 2000). It therefore appears that the main part of the Cadomian arc and its Avalonian counterpart are preserved in the Avalonia microcontinent (Murphy et al., 2006), whereas the main part of the back-arc basins remained in Cadomia. This arrangement is important to the subsequent opening of the Rheic Ocean (see below). Figure 13 shows a possible reconstruction of the Cadomian back-arc basin in the Saxo-Thuringian zone, with deposition of the passive margin sequences of the Weesenstein and Clanzschwitz
groups on the southern flank. The Rothstein and the Altenfeld formations are located in the interior portion of the back-arc basin within the external domain of the Saxo-Thuringian zone to the north (Fig. 2). We interpret the present geographical arrangement as reflecting the original paleoposition on the west African margin. Sample Pur-1 (Weesenstein Group) from the passive margin sequence contains ~40% Late Neoproterozoic detrital zircons derived from the Cadomian arc, which is considerably less than their abundance in Roth-1 from the Rothstein Formation (66%). In addition, the two samples differ in their Neoproterozoic age spectra, ~60% of Late Neoproterozoic grains in Pur-1 falling in the age range 670–600 Ma, whereas ~60% in Roth-1 fall in the range ca. 590–560 Ma. This difference suggests that in the source area of the passive margin, sequence remnants of an earlier stage of the Cadomian magmatic arc were exposed. This arc may represent the transitional zone between the craton and the stretched crust underlying the basin floor of the back-arc basin (Fig. 13). Consequently, we assume this early arc stage (ca. 670–600 Ma) was characterized by a Cordilleran-type active margin with subduction of the oceanic plate directly under the craton. The Cadomian back-arc basin probably opened into an expanded marginal basin to allow for the differentiation of its deposits into various facies patterns. The age of granitoid pebbles (577 ± 3 and 568 ± 4 Ma; Linnemann et al., 2000) from the Weesenstein and Clanzschwitz groups suggests rapid exhumation of granitoids intruded in the source area during or shortly before opening of the back-arc basin. The latter likely took place within a strike-slip regime, with the development of small subbasins acting as local suppliers of sediments, including material derived from the underlying stretched crust (Fig. 13). On the basis of similar relationships in Avalonia, Nance and Murphy (1994) proposed a model in which the oblique vector of subduction beneath the arc led to strike-slip motions in the back-arc basin. The concept of oblique subduction and its effects in the hinterland, combined with our field observations, is incorporated into the model in Figure 13. Cadomian Retroarc Basin The largest part of the weakly metamorphosed and well-preserved Late Neoproterozoic sediments of the Saxo-Thuringian zone is represented by the Lausitz Group (Wett-1; Lausitz antiform), the Leipzig Formation (North Saxon antiform), and the Frohnberg Formation (southeastern part of the Schwarzburg antiform) (Figs. 2 and 3). These units are positioned between the deposits of the inner back-arc basin and the passive margin (Figs. 2 and 3) and clearly differ in their flysch-like character from the sedimentary units discussed above. They are characterized by monotonous series of dark-gray graywacke turbidite intercalations of conglomerates and microconglomerates that often contain fragments of granitoids, metasediments, and black cherts (Fig. 4D). Frequent black chert fragments indicate their derivation, in large part, from eroded material derived from the inner back-arc basin. Given their similarity in lithology, sedimentation
The continuum between Cadomian orogenesis and opening of the Rheic Ocean
Cadomian retroarc basin c. 545-540 Ma
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Lausitz Group (Wett-1: 543+/-2 Ma) +Leipzig fm. Frohnberg Magmatic and Formation anatectic event (lower at c. 540 Ma section) Change to a transform margin
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Figure 14. Model for the plate-tectonic evolution of the Cadomian retroarc basin between ca. 545 and 540 Ma, based on data from the SaxoThuringian zone. There is no Cadomian angular unconformity on the continentward outer margin of the retroarc basin because Late Neoproterozoic sediments were unaffected by deformation (e.g., upper section of Frohnberg Formation in southeastern part of Schwarzburg antiform). In contrast, closer to the fold-and-thrust belt in the inner part of the retroarc basin, the sediments are deformed and consequently an angular unconformity is developed between Neoproterozoic retroarc sediments and overstepping Cambro-Ordovician strata (e.g., Cadomian unconformities at top of the Lausitz Group and Leipzig Formation). fm—formation.
regime, spatial distribution, and detrital modes, a comparable depositional age is assumed for the Leipzig and Frohnberg formations and the Lausitz Group. Depostion most likely took place at the Precambrian-Cambrian boundary, ~20–30 m.y. after the opening of the back-arc basin, as inferred from Wett-1 zircon ages and the intrusion of younger granites (see above). Sedimentary structures and paleoseismic features, such as water-escape structures (Fig. 4E), soft pebbles, and seismites, suggest rapid sedimentation (Linnemann, 2007). As shown in Figure 14, we interpret the Lausitz Group and the Leipzig/Frohnberg formations to be parts of a Cadomian retroarc basin or foreland basin. In our model this basin was formed during closure of the back-arc basin in response to the collision of the Cadomian
arc with the West African craton. Only the inner part of the retroarc basin was folded and thrusted, which explains why the more northerly portions of the Lausitz Group and the Leipzig Formation were deformed before intrusion of the ca. 540-Ma granitoids (Linnemann et al., 2000). An angular Cadomian unconformity between Late Neoproterozoic rocks and Cambro-Ordovician deposits is documented from different sections of these units (Linnemann and Buschmann, 1995a,b). However, the zircon age of 486 ± 4 Ma from pyroclastic sediments (KArc-1) of the Frohnberg Formation point to an Early Ordovician onset of Paleozoic sedimentation in the sections from the southeastern part of the Schwarzburg antiform. Cambrian strata are missing here, and the Cadomian unconformity is not an angular one. Instead, it is a
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disconformity or a simple sedimentation gap (paraconformity), without the occurrence of deformation between the under- and overlying strata. Thus, the undeformed Frohnberg Formation is interpreted to have been deposited in a more distal position relative to the fold-and-thrust belt and more proximal to the cratonic hinterland, whereas the deformed Lausitz Group and Leipzig Formation were situated closer to the colliding arc (Fig. 14). In contrast to the lower part of the Frohnberg Formation, which is composed mainly of thick-bedded graywackes, the ~100-m-thick upper part was likely deposited in a shallower marine environment (Linnemann, 2003b). This observation suggests termination of the retroarc basin regime and the onset of deposition in a remnant basin located in front of the outer retroarc deposits (Fig. 14). The very topmost quartzite bed of the section, known as “Basisquarzit” (= basal quartzite), is tradionally interpreted as the base of the Paleozoic overstep sequence (von Gaertner, 1944). It is, however, more likely that this quartzite represents the final bed of a continuous upward-thickening section of the Neoproterozoic remnant basin. The Cadomian retroarc basin and the related remnant basin were short-lived depositional systems. Based on the youngest detrital zircon ages of Wett-1 (543 ± 4 Ma; Fig. 11A) and the age of the Lausitz granitoid complex (539 ± 6 Ma; Fig. 3), which intrudes these sedimentary units, the ages and time spans of these systems can be confined to a time interval of <12 m.y. at the end of the Neoproterozoic and in the earliest Cambrian. Most Late Neoproterozoic sedimentary were intruded by voluminous granitoids at ca. 540 Ma (Linnemann et al., 2000; Gehmlich, 2003; Tichomirowa, 2003; Fig. 3). The largest exposed body is the Lausitz granitoid complex, which covers an area of ~100 × 50 km2. For this complex, a minimum granitoid volume of ~5000 km3 can be calculated, assuming a thickness of only 1 km. Most granitoids were likely derived from melting of the Late Neoproterozoic graywackes or similar units, because they contain large numbers of inherited zircons with age spectra comparable to those in the sediments (Linnemann et al., 2000, 2004; Gehmlich, 2003; Tichomirowa, 2003). This interpretation is supported by geochemical data and graywacke xenoliths in the granitoids (Hammer, 1996). These ca. 540-Ma granitoids record a relatively short-lived regime of high levels of heat flow, which we attribute in our model to slab break-off of the subducted oceanic plate (Fig. 14). All the processes summarized in Figures 13 and 14, from the early stages of a Cadomian magmatic arc (ca. 650–600 Ma), through opening of the Cadomian back-arc basin and its closure during arc-continent collision with subsequent formation of a retroarc basin, to the magmatic-anatectic event at ca. 540 Ma, correspond to our present understanding to the Cadomian orogen that formed at the margin of the West African craton. Opening of the Rheic Ocean There is no sharp break between the geological history linked to the Cadomian orogen and that of the Cambro-Ordovician, which finally led to the opening of the Rheic Ocean. Instead,
the latter is viewed as a logical continuation of the geological history of the dying marginal orogen. Nance and Murphy (1996) and Nance et al. (2002) proposed a Cordilleran model for the final stages of the Avalonian-Cadomian orogen analogous to the Cenozoic history of ridge-continent collision in the area of Baja California in the eastern Pacific. Such a model would explain both the geodynamic change from subduction-related processes to the opening of a new ocean and the excision of a long slice of continental crust, like that which formed the microcontinent of Avalonia. We have adapted these ideas to explain the platetectonic setting of the Saxo-Thuringian zone during the CambroOrdovician. After ridge-continent collision, slab break-off was triggered at ca. 540 Ma by the switch from an active margin to a transform margin setting (Fig. 15, inset). Cambrian sediments in the Saxo-Thuringian zone are restricted to the Lower and Middle Cambrian, with the onset of sedimentation occurring in the Atdabanian at ca. 530 Ma (Elicki, 1997). These units are characterized by carbonates with archaeocyatha, siliciclastic sediments, and red beds. The last were likely derived from erosion of laterite horizons generated on the denuded Cadomian orogen and the cratonic hinterland at ca. 540–530 Ma (Linnemann and Romer, 2002). These occurrences suggest a general uplift of the Cadomian orogen that was probably due to the rapid changes in plate-tectonic setting. In addition, the laterites and the occurrence of archaeocyatha point to deposition at low paleolatitudes. The overall change of the plate-tectonic regime is reflected by the onset of Cambrian sedimentation. Detritus of the Cambrian deposits was predominantly (~80%) derived from the Cadomian orogen, as inferred from the age spectrum of Kam1. The plate-tectonic setting may therefore have been similar to that of the present-day Basin and Range Province lying close to Baja California and the San Andreas fault (Fig. 15). In this way, stretching and thinning of the Cadomian crust and transcurrent faulting induced by the activity of the transform margin may have led to the opening of a rift basin filled with Lower and Middle Cambrian sediments. As a result, the Cadomian orogen became largely denuded. The rift basin likely developed on the side of the faulted and thrusted orogen, because this location would have been more sensitive to tectonic reactivation than the cratonic hinterland. The interplay between the more stable cratonic hinterland and the weaker Cadomian crust is thought to have led to asymmetric rifting (Fig. 15). Upper Cambrian sediments are relatively scarce in the Saxo-Thuringian zone, and fossiliferous deposits of this age are unknown. However, the lower and upper part of the Rollkopf Formation from the Vesser complex were deposited during the Middle and Upper Cambrian, respectively (see above). This unit belongs to the external domain of the Saxo-Thuringian zone, reflecting a paleoposition on the outer margin of the eroded and recycled Cadomian orogen. The Vesser complex is dominated by MOR-related igneous rocks associated with metasediments (Bankwitz et al., 1992; Figs. 2 and 6). In our view, this complex records the incision of an oceanic ridge that collided with the
The continuum between Cadomian orogenesis and opening of the Rheic Ocean
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Figure 15. Model for the formation of the asymmetric rift basin during the Lower to Middle Cambrian between ca. 530 and 500 Ma in the SaxoThuringian zone. The geological setting (see inset) is assumed to be similar to that of the Basin and Range province of North America 10 m.y. ago. Modified from Atwater (1970), Christiansen and Lipman (1972), Dickinson (1981), Condie (1989), and Nance et al. (2002).
periphery of the Cadomian orogen in a situation similar to present-day Baja California (inset of Fig. 16). Thinning of the lithosphere and upwelling asthenosphere led to enhanced heat flow in the upper lithosphere and the generation of Vesser magmatism. In our model the Vesser complex formed between the outer and inner zone of the asymmetric rift basin, as this location represents its weakest part (Fig. 16). The outer part, the former continental arc, was characterized by relatively thick crust, ongoing subsidence, and Upper Cambrian sedimentation, whereas the inner part, the former Cadomian back-arc basin, was more strongly affected by lithospheric thinning caused by uplift and upwelling of the asthenosphere (Fig. 16). Asymmetric rifting typically shows uplift of the remaining, thinned, lower plate and subsidence of the departing heavier upper plate (Wernicke, 1985; Coward, 1986). This asymmetry would explain the ongoing Cambrian sedimen-
tation on the upper plate, which would later become a part of Avalonia or a related terrane, and the absence of Upper Cambrian deposits on the lower plate, which represents the Cadomian realm at the periphery of the West African craton. This scenario is in agreement with the lack of Upper Cambrian sediments in the Saxo-Thuringian zone and the high maturity of Lower Ordovician deposits resulting from intense chemical weathering processes during the Upper Cambrian. Lower Ordovician deposits in the Cadomian part of central and western Europe are characterized by thick and widespread sandstone deposits, frequently metamorphosed to quartzites. The most prominent example is the ≤700-m-thick Armorican quartzite of the Armorican and Iberian massifs. Its equivalents in the Saxo-Thuringian zone are the quartzites of the 3000-m-thick Frauenbach and Phycodes groups from the southeastern part
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Vesser Uplift, chemical weathering, complex: and gap during bimodal the Upper Cambrian magmatism & MOR- Inner Zone Continent related of the (cratonic crust) rift basin rocks rift basin Subsidence (Ves-1: and ongoing 497+/-2 Ma) Upper Cambrian sedimentation
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Figure 16. Plate-tectonic model for the opening of the Rheic Ocean during the Upper Cambrian between ca. 500 and 490 Ma in the SaxoThuringian zone. Ocean opening is assumed to have been caused by the oblique subduction of an oceanic ridge similar to the present platetectonic situation (see inset) on the west coast of North America. MOR—mid-oceanic ridge. Modified from Atwater (1970), Christiansen and Lipman (1972), Dickinson (1981), Condie (1989), and Nance et al. (2002).
of the Schwarzburg antiform (Fig. 6). Quartzites of the Langer Berg Formation (Lbq-1) and those from the Hundsrück Group (Vesser complex) are their stratigraphic equivalents (Fig. 6). These deposits overstep in places Lower to Middle Cambrian strata (e.g., drill holes Heinersdorf 1 and 2; Wucher, 1967; Linnemann et al., 2000) and in other cases overlie directly the Cadomian basement, as in the Hohe Dubrau area in the Lausitz antiform (Linnemann and Buschmann, 1995b). Both forms of the Cadomian unconformity are also reported from the Armorican Massif and different parts of Iberia. It is therefore likely that rifting culminated in the Upper Cambrian, with the formation of rift shoulders, tilted blocks, and/or horsts and grabens, such that, in some places the Lower to Middle Cambrian is preserved, whereas in others the underlying Cadomian basement is exposed.
Extension over the entire paleolandscape and the enormous thickness of the Lower Ordovician overstep sequences classify these siliciclastics as post-rift sediments or deposits of a rift-drift transition. These sedimentary rocks must have been linked to considerable tectonic activity and thermal subsidence, resulting in large systems of detachment faults and escarpments on the surface. Lbq-1 (Langer Berg Formation) shows a very similar preOrdovician zircon age spectrum to those of the Neoproterozoic deposits (Table 4). The significant number of Paleoproterozoic and Archean grains suggests that rift-related breakaway faults extended into the cratonic hinterland (Fig. 17). In addition, Cadomian basement and Cambrian igneous rocks were either available to erosion or their zircons were distributed in the weathering crust formed during the Upper Cambrian and recycled by the Lower Ordovician transgression.
The continuum between Cadomian orogenesis and opening of the Rheic Ocean
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Final rift-related Remnants of Lower magmatism to Middle Cambrian at c. 490 Ma
Rift and drift transition: Opening of Rheic ocean and formation c. 490-480 Ma of passive margin Figure 17. Plate-tectonic model for the final opening of the Rheic Ocean and the formation of passive margins at ca. 490–480 Ma. The northern terrane that separated from the Gondwanan margin (Cadomia) could be part of Avalonia or a correlative terrane. Note the different unconformities between Neoproterozoic/Cambrian, Neoproterozoic/Ordovician and Cambrian/Ordovician strata, the general overstep of Lower Ordovician shallow marine sediments, the burial of the Vesser complex, and the renewed exhumation of cratonic crust in the hinterland. Inset shows analogous plate tectonic situation in part of the Basin and Range province of North America at ca. 3 Ma. Modified from Atwater (1970), Christiansen and Lipman (1972), Dickinson (1981), Condie (1989), and Nance et al. (2002).
Magmatic rocks with an age of ca. 490–480 Ma are widely distributed in the Saxo-Thuringian zone. They range from large plutons (Rumburk granite, Lausitz antiform; Figs. 2 and 3) to subvolcanic porphyroids (Bärentiegel, Schwarzburg antiform; Fig. 3) and various pyroclastics (e.g., Wurzelberg tuffite; Linnemann et al., 2000). An example of the last is sample KArc1 (“Konglomeratische Arkose”), dated at 486 ± 4 Ma (Figs. 3, 11A). This magmatic event represents the final rift-related magmatism (Fig. 17). After ca. 480 Ma, the Saxo-Thuringian zone is characterized by tectonic and magmatic quiescence and monotonous shelf sedimentation. We interpret this zone to be the passive margin of the Rheic Ocean, which had opened as a result of the separation of Avalonia or a related terrane (Fig. 18).
The geodynamic evolution outlined above has been deduced as far as possible from field geology, zircon dating, and petrographic and geochemical data. We have attempted to illustrate the way Cadomian orogenic processes may have operated and have compiled evidence supporting the idea that large parts of the Cadomian magmatic arc are now present in Avalonia or a related terrane, while the back-arc and the retroarc basins remained in Cadomia. We believe the long-lasting subduction processes that produced the arc and related basins were terminated by ocean closure in combination with ridge-continent collision. We have also tried to show that Cadomian oroenic processes and the subsequent opening of the Rheic Ocean were closely related to each other.
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c. 480 Ma
into the continent caused the formation of rift basins during the Lower to Middle Cambrian (530–500 Ma). This plate-tectonic scenario is assumed to have developed in a setting similar to that of the Baja California area ca. 3–10 Ma ago. This process continued from the Middle to the Upper Cambrian (ca. 500–490 Ma) and finally caused the opening of the Rheic Ocean, an event documented by thick Lower Ordovician siliciclastic sediments and a final rift-related, bimodal magmatic event at ca. 490–480 Ma.
e th of an g in ce en ic o a p e O h domi R Ca B
ACKNOWLEDGMENTS
Panthalassa ocean Equator
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Figure 18. Paleogeography in the Lower Ordovician during the opening of the Rheic Ocean at ca. 480 Ma. A—Armorican Massif and French Massif Central; B—Teplá-Barrandian unit (Bohemian Massif); I—Iberia; SX—Saxo-Thuringian zone (Bohemian Massif). Modified after C.R. Scotese (paleomap web site: www.scotese.com).
CONCLUSIONS Sediment provenances and magmatic events of Late Neoproterozoic (Ediacaran) and Cambro-Ordovician rock assemblages of the Saxo-Thuringian zone have been constrained by new LAICP-MS U-Pb ages from detrital zircons of five sandstones and magmatic zircons in an ignimbrite and one tuffite. These geochronological results, in combination with the analysis of platetectonic setting constrained from field observations, sedimentological and geochemical data, and trends in basin development, have been used to reconstruct Cadomian orogenic processes during the Late Neoproterozoic and the earliest Cambrian. A continuum between Cadomian orogenesis and the opening of the Rheic Ocean in the Cambro-Ordovician is supported by this data set. The early stage of Cadomian evolution is characterized by a Cordilleran-type continental magmatic arc, which was established at the periphery of the West African craton between 650 and 600 Ma. Subsequently, at ca. 590–560 Ma, a back-arc basin was formed behind the Cadomian magmatic arc. The formation of this basin was caused by crustal stretching in a strike-slip regime, which is similar to that presently observed in the Japan Sea of the western Pacific. Following collision of the Cadomian magmatic arc with the cratonic hinterland, the back-arc basin was closed between ca. 545 and 540 Ma. At this time a short-lived Cadomian retroarc basin was formed. Subsequently, a mid-oceanic ridge was subducted underneath the Cadomian orogen. This process may have been accompanied by slab break-off of the subducted oceanic plate, which resulted in increased heat flow, reflected in voluminous magmatic and anatectic events culminating at ca. 540 Ma. The subsequent oblique incision of the oceanic ridge
Prof. Dr. Gerhard Brey, Dr. Heidi Höfer, Jan Heliosch, Kai Klama, and Dr. Yann Lahaye (Institut für Geowissenschaften, Facheinheit Mineralogie, Universität Frankfurt am Main) are thanked for their help and fruitful discussions. We thank Damian Nance (Athens, Ohio, United States) for editorial handling and improving the language, and Dr. Jana M. Horák (Cardiff, United Kingdom), Dr. Stanislaw Mazur (Wroclaw, Poland) and Dr. Armin Zeh (Würzburg, Germany) for their critical reviews and inspiring discussions. Funding was provided to AG by the German Science Foundation (DFG; GE 1152/2-2). This paper is a contribution to the International Geological Correlation Program Project 497—“The Rheic Ocean: Its origin, evolution and correlatives” (http://www. snsd.de/igcp497/). REFERENCES CITED Atwater, T., 1970, Implications of plate tectonics for the Cenozoic tectonic evolution of western North America: Geological Society of America Bulletin, v. 81, p. 3513–3536. Avigad, D., Kolodner, K., McWiliams, M., Persing, H., and Weissbrod, T., 2003, Origin of northern Gondwana Cambrian sandstone revealed by detrital zircon SHRIMP dating: Geology, v. 31, p. 227–230, doi: 10.1130/00917613(2003)031<0227:OONGCS>2.0.CO;2. Bankwitz, P., Bankwitz, E., Kramer, W., and Pin, C., 1992, Early Paleozoic bimodal volcanism in the Vesser area, Thuringian Forest, eastern Germany: Zentralblatt für Geologie und Paläontologie, Teil I, v. 9/10, p. 1113–1132. Barrois, C., 1899, Sketch of the geology of Brittany: Proceedings of the Geologists’ Association. Geologists’ Association, v. 16, p. 101–132. Bertrand, L., 1921, Historie de la formation du sous-sol de la France. 1. Les anciennes mers de la France et leurs dépôts: Paris, Ernest Flammarion, 190 p. Bowring, S.A., Ramezani, J., and Grotzinger, J.P., 2003, High-precision U-Pb geochronology and the Cambrian-Precambrian boundary, in Geological Association of Canada, NUNA Conference 2003, New frontiers in the fourth dimension: Generation, calibration and application of geological time scales, 15–18 March, Mont Tremblant, Quebec. Bunel, H., 1835, Observations sur les terrains intermédiaires du Calvados: Mémoires de la Société Linnéenne de Normandie, v. 5, p. 99–100. Buschmann, B., 1995, Geotectonic facies analysis of the Rothstein Formation (Neoproterozoic, Saxothuringian Zone, east Germany) [Ph.D. thesis]: Freiberg, TU Bergakademie Freiberg, 122 p. Buschmann, B., Nasdala, L., Jonas, P., Linnemann, U., and Gehmlich, M., 2001, SHRIMP U-Pb dating of tuff-derived and detrital zircons from Cadomian marginal basin fragments (Neoproterozoic) in the northeastern Saxothuringian Zone (Germany): Neues Jahrbuch Geologie Paläontologie, Monatshefte, v. 2001, p. 321–342. Buschmann, B., Elicki, O., and Jonas, P., 2006, The Cadomian unconformity in the Saxo-Thuringian Zone, Germany: Palaeogeographic affinities of Ediacaran (terminal Neoproterozoic) and Cambrian strata: Precambrian Research, v. 147, no. 3–4, p. 387–403, doi: 10.1016/j.precamres.2006.01.023. Christiansen, R.L., and Lipman, P.W., 1972, Cenozoic volcanism and plate-tectonic evolution of the western United States: Part 2, Late Cenozoic: Royal Society of London Philosophical Transactions, ser. A, v. 271, p. 249–284.
The continuum between Cadomian orogenesis and opening of the Rheic Ocean Condie, K.C., 1989, Plate tectonics and crustal evolution: Oxford, Pergamon Press, 492 p. Coward, M.P., 1986, Heterogeneous stretching, simple shear and basin development: Earth and Planetary Science Letters, v. 80, p. 325–336, doi: 10.1016/0012-821X(86)90114-7. Dickinson, W.R., 1981, Plate tectonics and the continental margin of California, in Ernst, W.G., ed., The geotectonic development of California: Englewood Cliffs, New Jersey, Prentice-Hall, p. 1–28. D’Lemos, R.S., Iglis, J.D., and Samson, S.D., 2006, A newly discovered orogenic event in Morocco: Neoproterozoic ages for supposed Eburnean basement of the Bou Azzer inlier, Anti-Atlas Mountains: Precambrian Research, v. 147, no. 1–2, p. 65–78, doi: 10.1016/j.precamres.2006.02.003. Drost, K., Linnemann, U., McNaughton, N., Fatka, O., Kraft, P., Gehmlich, M., Tonk, C., and Marek, J., 2004, New data on the Neoproterozoic-Cambrian geotectonic setting of the Teplá-Barrandian volcano-sedimentary successions: Geochemistry, U-Pb zircon ages, and provenance (Bohemian Massif, Czech Republic): International Journal of Earth Sciences, v. 93, p. 742–757, doi: 10.1007/s00531-004-0416-5. Dufrenoy, P.A., 1838, Mémoire sur l’áge et la composition des terrains de transition de l’Ouest de la France: Annales des Mines, v. 3, p. 91–100. Elicki, O., 1997, Biostratigraphic data of the German Cambrian—present state of knowledge: Freiberger Forschungshefte, v. C 466, p. 155–165. Fernández-Suárez, J., Gutiérrez-Alonso, G., and Jeffries, T.E., 2002, The importance of along-margin terrane transport in northern Gondwana: Insights from detrital zircon parentage in Neoproterozoic rocks from Iberia and Brittany: Earth and Planetary Science Letters, v. 204, p. 75–88, doi: 10.1016/S0012-821X(02)00963-9. Finger, F., Hanzl, P., Pin, C., von Quadt, A., and Steyrer, H.P., 2000, The Brunovistulian: Avalonian Precambrian sequence at the eastern end of the central European Variscides, in Franke, W., Haak, U., Oncken, O., and Tanner, D., eds., Orogenic processes: Quantification and modelling in the Variscan belt: London, Geological Society of London, Special Publication 179, p. 103–112. Friedl, G., Finger, F., McNaughton, N.J., and Fletcher, I.R., 2000, Deducing the ancestry of terranes: SHRIMP evidence for South America-derived Gondwana fragments in central Europe: Geology, v. 28, p. 1035–1038. Friedl, G., Finger, F., Paquette, J.L., von Quadt, A., McNaughton, N.J., and Fletcher, I.R., 2004, Pre-Variscan geological events in the Austrian part of the Bohemian Massif deduced from U-Pb zircon ages: International Journal of Earth Sciences, v. 93, p. 802–823. Gaertner, von, H.R., 1944, Die Schichtgliederung der Phyllitgebiete in Thüringen und Nordbayern und ihre Einordnung in das stratigraphische Schema: Jahrbuch der Reichsanstalt für Bodenforschung, v. 62, p. 54–80. Gebauer, D., and Friedl, G., 1994, A 1.38 Ga protolith age for the Dobra orthogneiss (Moldanubian zone of the Southern Bohemian Massif, NEAustria): Evidence from ion-microprobe (SHRIMP) dating of zircon: Journal of the Czech Geological Society, v. 39, no. 1, p. 34–35. Gebauer, D., Williams, I.S., Compston, W., and Grünenfelder, M., 1989, The development of the central European continental crust since the Early Archaean based on conventional and ion-microprobe dating of up to 3.84 b.y. old detrital zircons: Tectonophysics, v. 157, p. 81–96, doi: 10.1016/0040-1951(89)90342-9. Gehmlich, M., 2003, Die Cadomiden und Varisziden des Saxothuringischen Terranes—Geochronologie magmatischer Ereignisse: Freiberger Forschungshefte, v. C 500, p. 1–129. Gerdes, A., and Zeh, A., 2006, Combined U-Pb and Hf isotope LA-(MC-)ICPMS analyses of detrital zircons: Comparison with SHRIMP and new constraints for the provenance and age of an Armorican metasediment in central Germany: Earth and Planetary Science Letters, v. 249, no. 1–2, p. 47–62, doi: 10.1016/j.epsl.2006.06.039. Glasmacher, U., Bauer, W., Clauer, N., and Puchkov, V.N., 2004, Neoproterozoic metamorphism and deformation at the southeastern margin of the East European craton, Uralides, Russia: International Journal of Earth Sciences, v. 93, p. 921–944, doi: 10.1007/s00531-004-0426-3. Gradstein, F., Ogg, J., and Smith, A., eds., 2004, A geologic time scale 2004: Cambridge, Cambridge University Press, 589 p. Graindor, M.J., 1957, Le Briovérien dans le Nord-Est du Massif armoricain: Mémoires pour servir a l’explication de la carte géologique détaillée de la France: Paris, Imprimerie Nationale, 211 p. Hammer, J., 1996, Geochemie und Petrogenese der cadomischen und spätvariszischen Granitoide der Lausitz: Freiberger Forschungshefte, v. C 463, p. 1–107. Hegner, E., and Kröner, A., 2000, Review of Nd isotopic data and xenocrystic and detrital zircon ages from the pre-Variscan basement in the eastern Bohemian Massif: Speculations on palinspastic reconstructions, in Franke, W., Haak, U., Oncken, O., and Tanner, D., eds., Orogenic pro-
95
cesses: Quantification and modelling in the Variscan belt. Geological Society of London, Special Publication, v. 179, p. 113–129. Heuse, T., Kroner, U., and Bartl, S., 2001, Tektono- und lithostratigraphische Kartierungskriterien für den Zentralteil des Schwarzburger Sattels: Geowissenschaftliche Mitteilungen von Thüringen, v. 9, p. 107–124. Hoskin, P.W.O., and Schaltegger, U., 2003, The composition of zircon and igneous and metamorphic petrogenesis: Reviews in Mineralogy and Geochemistry, v. 53, p. 27–62, doi: 10.2113/0530027. Janoušek, V., Gerdes, A., Vrána, S., Finger, F., Erban, V., Friedl, G., and Braithwaite, C.J.R., 2006, Low-pressure Granulites of the Lišov Massif, southern Bohemia: Viséan metamorphism of Late Devonian plutonic arc rocks: Journal of Petrology, v. 47, p. 705–744, doi: 10.1093/petrology/egi091. Jolivet, L.M., Fournier, P., Huchnon, V.S., Rozhdestvenskiy, K.F.S., and Oscorbin, L.S., 1992, Cenozoic intracontinental dextral motion in the Okhotsk–Japan Sea region: Tectonics, v. 11, p. 968–977. Kemnitz, H., Romer, R.L., and Oncken, O., 2002, Gondwana break-up and the northern margin of the Saxothuringian belt (Variscides of central Europe): International Journal of Earth Sciences, v. 91, p. 246–259, doi: 10.1007/ s005310100209. Keppie, J.D., Davies, D.W., and Krogh, T.E., 1998, U-Pb geochronological constraints on Precambian stratified units in the Avalon composite terrane of Nova Scotia, Canada: Tectonic implications: Canadian Journal of Earth Sciences, v. 35, p. 222–236, doi: 10.1139/cjes-35-3-222. Kerforne, F., 1901, Discordance du Cambrien sur le Precamrien, près de Rennes: Bulletin de la Société Géologique de France, v. 4, no. I, p. 2587–2590. Kossmat, F., 1927, Gliederung des varistischen Gebirgsbaues: Abhandlungen des Sächsischen Geologischen Landesamtes, v. 1, p. 1–39. Krˇíbek, B., Pouba, Z., Skocˇek, V., and Waldhausrová, J., 2000, Neoproterozoic of the Teplá-Barrandian unit as a part of the Cadomian orogenic belt: A review and correlation aspects: Veˇstnik Cˇ eskeho geologického ústavu, v. 75, no. 3, p. 175–196. Kroner, U., Linnemann, U., and Romer, R.L., 2003, Synthese der geologischen Geschichte des Saxothuringikums: Vom cadomischen Akkretionsorogen zum variszischen Kollisionsgebirge, in Linnemann, U., ed., Das Saxothuringikum: Dresden/Bautzen, Staatliche Naturhistorische Sammlungen Dresden, p. 147–150. Kroner, U., Hahn, T., Romer, R.L., and Linnemann, U., 2007 (this volume), The Variscan orogeny in the Saxo-Thuringian zone—Heterogenous overprint of Cadomian/Paleozoic Peri-Gondwana Crust, 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: Boulder, Colorado, Geological Society of America Special Paper 423, doi: 10.1130/2007.2423(06). Linnemann, U., 1991, Glazioeustatisch kontrollierte Sedimentationsprozesse im Oberen Proterozoikum der Elbezone (Weesensteiner Gruppe/Sachsen): Zentralblatt für Geologie und Paläontologie, v. 1, no. 12, p. 2907–2934. Linnemann, U., 1996, Die Lagerungsverhältnisse und die sedimentäre Entwicklung des basalen Altpaläozoikums vom Langen Berg bei Gehren (Schwarzburger Antiklinorium, Saxothuringikum): Beiträge zur Geologie von Thüringen, v. 3, p. 73–84. Linnemann, U., 2003a, The structural units of Saxo-Thuringia: Geologica Saxonica, v. 48/49, p. 19–28. Linnemann, U., 2003b, Sedimentation and geotectonic setting of the basin development of Saxo-Thuringia (Neoproterozoic–Lower Carboniferous): Geologica Saxonica, v. 48/49, p. 71–110. Linnemann, U., 2007, Ediacaran rock complexes, the Cadomian orogeny and the problem of the Precambrian-Cambrian boundary in Germany (NE Bohemian Massif, central Europe): London, Geological Society of London Special Publication (in press). Linnemann, U., and Buschmann, B., 1995a, Der Nachweis der cadomischen Diskordanz in einer Tiefenbohrung bei Gera und deren Bedeutung für das proterozoisch-paläozoische Standardprofil im Schwarzburger Antiklinorium: Geowissenschaftliche Mitteilungen von Thüringen, v. 3, p. 1–11. Linnemann, U., and Buschmann, B., 1995b, Die cadomische Diskordanz im Saxothuringikum (oberkambrisch-tremadocische overlap-Sequenzen): Zeitschrift für Geologische Wissenschaften, v. 23, no. 5/6, p. 707–727. Linnemann, U., and Heuse, T., 2000, The Ordovician of the Schwarzburg anticline: Geotectonic setting, biostratigraphy and sequence stratigraphy (Saxo-Thuringian Terrane, Germany): Zeitschrift der deutschen geologischen Gesellschaft, v. 151, p. 471–491. Linnemann, U., and Romer, R.L., 2002, The Cadomian orogeny in SaxoThuringia, Germany: Geochemical and Nd-Sr-Pb isotopic characterization of marginal basins with constraints to geotectonic setting and provenance: Tectonophysics, v. 352, p. 33–64, doi: 10.1016/S0040-1951(02)00188-9.
96
Linnemann et al.
Linnemann, U., and Schauer, M., 1999, Die Entstehung der Elbezone vor dem Hintergrund der cadomischen und variszischen Geschichte des Saxothuringischen Terranes—Konsequenzen aus einer abgedeckten geologischen Karte: Zeitschrift für Geologische Wissenschaften, v. 27, p. 529–561. Linnemann, U., Gehmlich, M., Tichomirowa, M., Buschmann, B., Nasdala, L., Jonas, P., Lützner, H., and Bombach, K., 2000, From Cadomian subduction to Early Paleozoic rifting: The evolution of Saxo-Thuringia at the margin of Gondwana in the light of single zircon geochronology and basin development (central European Variscides, Germany): London, Geological Society of London Special Publication 179, p. 131–153. Linnemann, U., McNaughton, N.J., Romer, R.L., Gehmlich, M., Drost, K., and Tonk, C., 2004, West African provenance for Saxo-Thuringia (Bohemian Massif): Did Armorica ever leave pre-Pangean Gondwana? U/Pb-SHRIMP zircon evidence and the Nd-isotopic record: International Journal of Earth Sciences, v. 93, p. 683–705, doi: 10.1007/s00531-004-0413-8. Ludwig, K.R., 2001, User’s manual for Isoplot/Ex rev. 2.49: Berkeley, Berkeley Geochronology Center Special Publication 1a, 56 p. Massonne, H.-J., 1998, A new occurrence of microdiamonds in quartzofeldspathic rocks of the Saxonian Erzgebirge, Germany, and their metamorphic evolution: 7th International Kimberlite Conference, Cape Town, April 1998, extended abstracts, p. 552–554. Mazur, S., Turniak, K., and Bröcker, M., 2004, Neoproterozoic and CambroOrdovician magmatism in the Variscan Klodzko Metamorphic complex (west Sudetes, Poland): New insights from U-Pb zircon dating: International Journal of Earth Sciences, v. 93, p. 758–772. McKerrow, W.S., Mac Niocaill, C., Ahlberg, P.E., Clayton, G., Cleal, C.J., and Eagar, R.M.C., 2000, The Late Palaeozoic relations between Gondwana and Laurussia, in Franke, W., Haak, U., Oncken, O., and Tanner, D., eds., Orogenic processes: Quantification and modelling in the Variscan belt: London, Geological Society of London Special Publication 179, p. 9–20. Murphy, J.B., and Nance, R.D., 1991, Supercontinent model for the contrasting character of Late Proterozoic orogenic belts: Geology, v. 19, p. 469–472, doi: 10.1130/0091-7613(1991)019<0469:SMFTCC>2.3.CO;2. Murphy, J.B., Strachan, R.A., Nance, R.D., Parker, K.D., and Fowler, M.B., 2000, Proto-Avalonia: A 1.2–1.0 Ga tectonothermal event and constraints for the evolution of Rodinia: Geology, v. 28, p. 1071–1074, doi: 10.1130/0091-7613(2000)028<1071:PAAGTE>2.3.CO;2. Murphy, J.B., Pisarevsky, S.A., Nance, R.D., and Keppie, J.D., 2004, Neoproterozoic–Early Paleozoic evolution of Peri-Gondwanan terranes: Implications for Laurentia-Gondwana connections: International Journal of Earth Sciences, v. 93, p. 659–682. Murphy, J.B., Gutiérrez Alonso, G., Nance, R.D., Fernández-Suárez, J., Keppie, J.D., Quesada, C., Strachan, R.A., and Dostal, J., 2006, Origin of the Rheic Ocean: Rifting along a Neoproterozoic suture? Geology, v. 34, p. 325–328, doi: 10.1130/G22068.1. Nance, R.D., and Murphy, J.B., 1994, Contrasting basement isotopic signatures and the palinspastic restoration of peripheral orogens: Example from the Neoproterozoic Avalonian-Cadomian belt: Geology, v. 22, p. 617–620. Nance, R.D., and Murphy, J.B., 1996, Basement isotopic signatures and Neoproterozoic paleogeography of Avalonian-Cadomian and related terranes in the circum-North Atlantic, in Nance, R.D., and Thompson, M.D., eds., Avalonian and related peri-Gondwanan terranes of the circum-North Atlantic: Boulder, Colorado, Geological Society of America Special Paper 304, p. 333–346. Nance, R. D., Murphy, J. B., and Keppie, J. D., 2002, A Cordilleran model for the evolution of Avalonia: Tectonophysics, v. 352, p. 1–21. Oncken, O., 1997, Transformation of a magmatic arc and an orogenic root during oblique collision and its consequences for the evolution of the European Variscides (mid-German crystalline rise): Geologische Rundschau, v. 86, no. 1, p. 2–20. Roberts, D., and Siedlecka, A., 2002, Timanian orogenic deformation along the northeastern margin of Baltica, northwest Russia and northeast Norway, and Avalonian-Cadomian connections: Tectonophysics, v. 352, p. 169– 184, doi: 10.1016/S0040-1951(02)00195-6. Samson, S.D., D’Lemos, R.S., Blichert Toft, J., and Vervoort, J.D., 2003, U-Pb geochronology and Hf-Nd isotope compositions of the oldest
Neoproterozoic crust within the Cadomian orogen: New evidence for a unique juvenile terrane: Earth and Planetary Science Letters, v. 208, p. 165–180, doi: 10.1016/S0012-821X(03)00045-1. Schneider Santos, J.O., Hartmann, L.A., Gaudette, H.E., Groves, D., McNaughton, N.J., and Fletcher, I.R., 2000, A new understanding of the provinces of the Amazon craton based on integration of field mapping and U-Pb and Sm-Nd geochronology: Gondwana Research, v. 3, p. 453–488, doi: 10.1016/S1342-937X(05)70755-3. Sircombe, K.N., 2004, AgeDisplay: An Excel workbook to evaluate and display univariate geochronological data using binned frequency histograms and probability density distributions: Computers & Geosciences, v. 30, p. 21–31. Stacey, J.S., and Kramers, J.D., 1975, Approximation of terrestrial lead isotope evolution by a two-stage model: Earth and Planetary Science Letters, v. 26, p. 207–221, doi: 10.1016/0012-821X(75)90088-6. Tait, J.A., Bachtadse, V., Franke, W., and Soffel, H.C., 1997, Geodynamic evolution of the European Variscan fold belt: Paleomagnetic and geological constraints: Geologische Rundschau, v. 86, p. 585–598, doi: 10.1007/ s005310050165. Tichomirowa, M., 2003, Die Gneise des Erzgebirges—hochmetamorphe Äquivalente von neoproterozoisch-frühpaläozoischen Grauwacken und Granitoiden der Cadomiden: Freiberger Forschungshefte, v. C 495, p. 1–222. Unrug, R., 1996, Geodynamic map of Gondwana Supercontinent assembly: Orléans, France, Bureau de Recherches Géologiques et Minéres Orléans. Van der Voo, R., 1979, Paleozoic assembly of Pangea: A new plate tectonic model for the Taconic, Caledonian and Hercynian orogenies: Eos (Transactions, American Geophysical Union), v. 60, p. 241. Wendt, J.I., Kröner, A., Fiala, J., and Todt, W., 1993, Evidence from zircon dating for existence of approximately 2.1 Ga old crystalline basement in southern Bohemia, Czech Republic: Geologische Rundschau, v. 82., p. 42–50. Wendt, J.I., Kröner, A., Fiala, J., and Todt, W., 1994, U-Pb zircon and Sm-Nd dating of Moldanubian HP/HT granulites from south Bohemia, Czech Republic: Journal of the Geological Society of London, v. 151, p. 83–90. Wernicke, B., 1985, Uniform-sense normal simple shear of the continental lithosphere: Canadian Journal of Earth Sciences, v. 22, p. 108–125. Windley, B.F., 1995, The evolving continents (3rd edition): Chichester, Wiley, 526 p. Wucher, K., 1967, Ergebnisse der Kartierungsbohrungen Heinersdorf 1/60 und 2/62 (Thüringisches Schiefergebirge): Jahrbuch für Geologie, v. 1, p. 297–323. Zeh, A., and Wunderlich, J., 2003, Mitteldeutsche Kristallinzone, in Seidel, G., ed., Geologie von Thüringen: Stuttgart, E. Schweitzerbart’sche Verlagsbuchhandlung, p. 24–51. Zeh, A., Brätz, H., Millar, I.L., and Williams, I.S., 2001, A combined zircon SHRIMP and Sm-Nd isotope study of high-grade paragneisses from the mid-German crystalline rise: Evidence for northern Gondwanan and Grenvillian provenance: Journal of the Geological Society of London, v. 158, p. 983–994. Zeh, A., Williams, I.S., Brätz, H., and Millar, I.L., 2003, Different age response of zircon and monazite during the tectono-metamorphic evolution of a high grade paragneiss from the Ruhla crystalline complex, central Germany: Contributions to Mineralogy and Petrology, v. 145, p. 691–706, doi: 10.1007/s00410-003-0462-1. Zeh, A., Gerdes, A., Will, T.M., and Millar, I.L., 2005, Provenance and magmatic-metamorphic evolution of a Variscan island-arc complex: Constraints from U-Pb dating, petrology, and geospeedometry of the Kyffhäuser crystalline complex, central Germany: Journal of Petrology, v. 46, p. 1393–1420, doi: 10.1093/petrology/egi020. 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 Gesellschaft für Geowissenschaften, v. 150, no. 4, p. 627–639. MANUSCRIPT ACCEPTED BY THE SOCIETY 3 OCTOBER 2006
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Geological Society of America Special Paper 423 2007
The Lausitz graywackes, Saxo-Thuringia, Germany— Witness to the Cadomian orogeny Helga Kemnitz* GeoForschungsZentrum Potsdam, Telegrafenberg, D-14473 Potsdam, Germany
ABSTRACT The Lausitz Group in the easternmost part of Saxo-Thuringia, Germany, forms the largest exposure of partly anchimetamorphic Cadomian basement in Germany. In common with adjoining units to the west, it was deposited in a convergent-margin basin of northern peri-Gondwana. Sedimentary features within the turbiditic graywacke successions suggest continuous accumulation during a deepening stage that followed basaltic-andesitic arc-volcanic activity. The turbidites are irregularly intercalated with tuffaceous graywacke, which was derived from reworked, but only slightly older, basic volcanic material. Petrological and geochemical data reveal a broadly uniform source area for the graywackes with a dissected magmatic arc signature containing exhumed pre-Cadomian basement. Graywackes of the Lausitz Group together with earlier Cadomian basin successions and parts of the pre-Cadomian basement became mobilized during Cadomian subduction-related anatexis. During basin closure, the graywackes were weakly folded with a northerly vergence. Subsequently, they were contact-metamorphic overprinted by granodioritic intrusions, which mark the end of the Cadomian orogeny and show both Cadomian and preCadomian crustal signatures. Superimposed very low-grade S-C textures in graywackes of the eastern Lausitz area are due to Variscan processes. Keywords: Lausitz Group, Cadomia, sediment, petrology, geochemistry, provenance analysis, lithostratigraphy INTRODUCTION
et al., 2004), (2) graywacke composition and convergent-margin continental arc signature, (3) partial melting and intrusion by Stype granitoids, (4) lack of high-pressure metamorphism, and (5) post-Cadomian Lower to Middle Cambrian transtensional history (Nance et al., 1991; Nance and Murphy, 1994). These successions, which now occur as Cadomian-type (Murphy et al., 2004) basement fragments within younger Phanerozoic orogenic realms, include the Teplá-Barrandian (Czech Republic), the Ossa-Morena Zone (southern Spain), the French Armorican Massif, units of the Moldanubian Zone, and the Lausitz–IzeraKarkonosze domain. The last represents the eastern basement
During the Neoproterozoic, an Andean-type orogenic belt developed marginal to north and northwest Gondwana. Thick successions of marine arc-related basin sediments accumulated within this belt (Nance and Murphy, 1994; Murphy et al., 2004) and were subsequently involved in Cadomian orogenic processes. Quite homogeneously, these sediments display a number of common features: (1) west African provenance (Linnemann *E-mail:
[email protected].
Kemnitz, H., 2007, The Lausitz graywackes, Saxo-Thuringia, Germany—Witness to the Cadomian orogeny, 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. 97–141, doi: 10.1130/2007.2423(04). For permission to copy, contact
[email protected]. ©2007 Geological Society of America. All rights reserved.
97
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Kemnitz
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TEPLÁBARRANDIAN
M
IAN NUB A LD ZONE O M Cadomian basement: a b
Schwarzwald
150 km
exposed covered
Vogesen
Figure 1. Simplified map of major structural elements of the Variscan orogen showing the distribution of Cadomian basement within the SaxoThuringian Zone (Germany) and its extension into the Izera-Karkonosze Mountains (Poland) and the Teplá-Barrandian (Czech Republic): 1— Lausitz Group, sedimentary Neoproterozoic; 2—Anatexites and granodiorites; 3—Weesenstein Group; 4—Outer and Inner Gray-Gneiss Group; 5—Clanzschwitz Group; 6—Rothstein Formation; 7—Katzhütte Group; 8—Leipzig Group (northwest Saxony). SA—Schwarzburg Anticline; MM—Münchberg Massif; GM—Granulite Massif; B—Berlin; C—Cottbus; D—Dresden; Gö—Görlitz; L—Leipzig; P—Prague; Kacz.—Gorie Kaczawskie. Exposed Cadomian basement comprises (a) sedimentary and (b) granitic and anatectic rocks.
continuation of the Saxo-Thuringian part of Armorica in Germany and Poland (Linnemann et al., 2000; Kemnitz et al., 2002). Cadomian Saxo-Thuringia faced the Avalonian-Cadomian orogenic belt from an external position (Linnemann et al., 2000). For decades, geological research in the study area was marked by a debate over whether the anatexites, granodiorites, and their associated sedimentary rocks were the product of a single orogenic cycle. This controversy was largely the result of contrasts in deformation features between the western Lausitz granodiorites and those to the east, which were thought to be of Precambrian age (Rimann, 1910; Ebert, 1943). Indeed, Ebert (1943) proposed the existence of at least two graywacke series of different ages. In contradiction, Möbus (1956) concluded that all the granodioritic intrusions, from the West Lausitz to the West Sudetes, were co-genetic and of very similar age, regardless of deformational overprint. It remained unclear, however, which—if any—of the intrusions and graywacke units were Variscan or older. According to von Gaertner (1964), however, Lausitz granodioritic intrusions (including those of the Czech Republic and Poland) not
only had a common source, but were also of Precambrian age, and the folding of the graywackes was a Precambrian event, too. The latter was proved first by Burmann (1966), who discovered a microflora (Favososphaera conglobata) of Vendian depositional age from graywacke outcrops and drill-holes west of the city of Cottbus (Fig. 1). In the following years, re-examination, supported by detrital component analysis and geochemical discrimination methods from the sediments (Kemnitz and Budzinski, 1991, 1994; Kemnitz, 1994, 1998), and subsequent geochronological data from the magmatic rocks— including Rb-Sr whole-rock ages, Pb-Pb model ages (Bielicki et al., 1989), single-zircon evaporation ages, U-Pb sensitive high-resolution ion microprobe (SHRIMP) zircon ages, and whole-rock Sr, Nd, and Pb model ages (Kröner et al., 1994; Linnemann et al., 2000; Tichomirowa et al., 2001; Linnemann and Romer, 2002)—led to a new understanding and acceptance of the position of the Lausitz area within the Cadomian terranes. The basic data and idea discussed here originated in a project that aimed to determine the lithostratigraphic position of the
The Lausitz graywackes, Saxo-Thuringia, Germany
L
+ + v v
Cadomian anatexitic rocks, subduction related Granodioritic intrusions, post-collisional Cadomian granodiorites with Variscan deformation Early Palaeozoic Rumburk-Izera granite, rift stage, with Variscan deformation Variscan granites and granodiorites (Carboniferous) Cenozoic basic volcanic rocks Phanerozoic cover
WE ST L
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Intercalations of Cadomian basic subduction-related rocks: type I, IIa
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Contact-metamorphic zonation of the Lausitz Group:
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CZECH REPUBLIC
Figure 2. Geological map of the Saxo-Thuringian Lausitz domain showing contact-metamorphic zonation and occurrences of tuffaceous graywackes and basic-rock enclaves. Roman numerals I–IV mark the centers of the study regions. KM—Kamenz; RB—Radeberg; GÖ—Görlitz; ZG—Zgorzelec.
sediments and their relationship to contact-metamorphic and anatectic processes. Supplemented by own new data, this work presents a compilation of all recent developments from sediment petrographical and/or petrological data. GEOLOGICAL OVERVIEW Occupying ~5,000 km2, the Neoproterozoic rocks in the Lausitz region (Lusatia), East Saxony, form the largest exposure of Cadomian basement in Germany. In northwestern Saxony, west of the Elbe Zone (Fig. 1), the Cadomian basement extends farther north under Phanerozoic cover sediments and is bound by the Mid-German Crystalline Zone. Most parts of the Lausitz region have undergone long-term, variable uplift since the Lower Cambrian. The initial uplift stage, which probably started during or immediately after the Cadomian collision, was characterized by high exhumation rates that led to the erosion of large parts of the contact-metamorphosed Neoproterozoic graywackes, as indicated by unconformably overlying Ordovician on-lap sediments (Linnemann and Buschmann, 1995). Thus, the deepest crustal exposures are centrally and southerly located and indicate abun-
dant synsedimentary melt production. These units are of undefined (Neo)Proterozoic stratigraphy and composed of anatectic rocks, with metabasic and contact metamorphosed graywacke rafts and a variety of S-type granodioritic intrusions. Sediment successions adjoin this zone to the north and northeast (Fig. 2). The Neoproterozoic sedimentary rocks consist of turbiditic graywackes ~2–3 km in thickness. They were deposited in a marine basin that underwent subduction-related processes during the Cadomian orogeny. This processing is unequivocally recorded in their sedimentological, geochemical, and structural features (Eidam, 1988; Kemnitz and Budzinski, 1991, 1994) and has been confirmed by geochronological, isotopic, and rare-earth geochemical data (Krauss et al., 1992, Hammer et al., 1999; Linnemann et al., 2000; Tichomirowa et al., 2001; Linnemann and Romer, 2002). The Lausitz graywackes are regarded as a single lithostratigraphic unit, the Lausitz Group (Kemnitz, 1994, 1997), which continues across the West Sudetes into Poland, where graywackes and S-type granodioritic intrusions of Cadomian ages (Borkowska et al., 1980; Krauss et al., 1992; Korytowski et al., 1993; Borkowska, 1994; Kemnitz and Budzinski, 1994; Z˙ elaz´niewicz et al., 2004) extend up to the Intra-Sudetic
Kemnitz
GENERAL CHARACTERISTICS OF THE LAUSITZ GROUP Lithology The graywacke successions can be described by a common reference facies scheme (Fig. 3) corresponding to the Bouma (1962) turbidite intervals Ta–e. Litho types and sedimentary characteristics of all stratigraphic sections investigated display a striking similarity. The thickness of the sequences ranges, on average, from decimeter to decameter scale. In larger outcrops, a gradual decrease in maximum grain size and thickness of the graywacke beds and an increase in pelite bed thicknesses toward the top are common. In this way, repeated sequences form megasequences with thicknesses of at least 40 m, starting with a coarse to fine gravel interval. A sequence typically begins with an intraclast horizon consisting of reworked, underlying pelitic material. In general, basal intervals (Ta; Fig. 3) comprise only medium- to coarse-grained graywackes, which constitute up to ~10% of a sequence, whereas
cm to dm
mm to dm
Grain size Mudstone
Sandstone Silt
fs
ms cs
fg
Te Td
v
v
v
tuffaceous greywacke (cm to dm)
Tb cm to m
Fault and partly (granodiorites) across it (Fig. 1). In contrast to the Cadomian granodioritic rocks, however, the Variscan amphibole-bearing quartzdioritic and monzonitic intrusions of the Lausitz region show significant I-type characteristics (Eidam et al., 1995; Hammer, 1996) as island-arc-like signatures originated from melting processes in a lower crust. Cadomian metamorphism did not exceed anchimetamorphic conditions; the degree of preservation of primary (sedimentary) structures, however, is dependent of the degree of both Cadomian postcollisional contact-metamorphism and Variscan metamorphic overprint. West of the Elbe Zone, comparable Neoproterozoic units locally experienced Variscan high-pressure/ high-temperature metamorphism (Mingram, 1998; Linnemann et al., 2000). East of this structure, the units show a generally east–west-directed propagation of the Variscan front that is indicated by kinematic markers as well as by the gradual decrease in metamorphic grade from the Izera-Karkonosze region to the eastern Lausitz area (Z˙ aba, 1985; Mazur and Kryza, 1996; Seston et al., 2000). As the Lausitz region was only marginally affected, the epizonal graywackes of this area are prime candidates for studies of Cadomian sedimentation. The first description of the Lausitz graywackes as rhythmical flysch sediments was that of Schwab (1962). A Vendian age for the Lausitz graywackes was first suggested by Burmann (1966) and has been confirmed on several occasions (Burmann, 1969, 1972, 1997; Lorenz and Burmann, 1972; Weber et al., 1990). Burmann described organic-walled microfossils (favosospheres) and algae filaments whose occurrences, whether as isolated microfossils or as assemblages, can be traced across the Neoproterozoic Saxo-Thuringian (Burmann, 2000). These occurrences can be correlated with the Upper Brioverian of Bohemia (Konzalova, 1974) and Brittany (Chauvel and Mansúy, 1981).
Scale of thickness:
100
Ta
Figure 3. Facies model for the Lausitz Group as medial turbiditic graywackes of Bouma (1962) intervals Ta–Te. Tuffaceous graywacke intercalations are randomly distributed.
>50% of the lithoclasts are restricted to the grain-size interval of 250–550 μm. The fine-gravel fraction is rarely present. The fine-sand graywacke interval (Tb) and the overlying fine-sand to siltstone intervals (Tc and Td) represent between 50 and 85% of all sections examined. Interval Tc, composed of graywackes alternating with silt- and mudstone laminae, is characterized by a number of mainly low to medium energy features. These comprise cross-bedding, load, and erosion marks (flute casts, ripple marks, bifurcated rill marks, furrow casts, parting lineations, and grooves) and convolute bedding as well as minor slump structures (c.f. Schöbel, 1985; Rathner, 2001). The uppermost pelitic sediments (interval Te) consist predominantly of medium to coarse silt and make up ~5–15% of the section examined. Their geochemistry differs from the element pattern of marine pelites (Kemnitz and Budzinski, 1994), which confirms their common origin with the psammitic material from turbiditic flows. An intermediate to basic tuffaceous graywacke (Kemnitz and Budzinski, 1991; Kemnitz, 1997) occurs as <1%, but as a characteristic intercalation. It forms layers and lens-shaped bodies 1–15 cm in thickness and is not bound to any particular interval (Fig. 3). In view of the grain size distribution and lack of proximal facies features, the Lausitz Group is best interpreted as a hemipelagic deposit of a medial to prograding distal fan (Walker, 1979).
The Lausitz graywackes, Saxo-Thuringia, Germany Sedimentary and Petrographic Features According to Füchtbauer (1988), the Lausitz graywackes are lithic graywackes, with a lithoclasts:matrix ratio of ~1:1. On average, unstable components make up >45% of the mode, 7–15% of which are rock fragments (Table 1). Magmatic source rocks make up 85–95% of all kinds of lithoclasts (Table 1). Five to twenty percent of lithoclasts show crypto- to microcrystalline recrystallization of a glassy matrix; porphyric textures with quartz and feldspar phenocrysts; and subordinate spherolithic, trachytic, and subophitic textures, indicating a mainly intermediate to acid volcanic source. In the group of unstable components, the average abundance of single plagioclase grains (P; albite to minor oligoclase; Table 1) exceeds that of potassium feldspar (K) in the majority of samples. In addition to intrusive source rocks, much of the albite may have been derived from trachytic to dacitic volcanic rocks, as suggested by the positive correlation between the amount of P and intermediate volcanic rock fragments. In some cases, volcanic fragments (Lv) amount to over 40% of lithoclasts. Intrusive magmatic rocks are represented by single quartz grains (QM), quartz-feldspar aggregates (QF), and probably by the majority of K. This content is further supported by granophyric or myrmekitic textures, which indicate granitic to granodioritic or tonalitic source rocks. Most QM show typical features of cataclastic deformation (including dynamic recrystallization) that is assumed to be a result of syn- to postintrusive deformation of the source rocks. Rock fragments of unambiguous metamorphic origin (Lm) represent 0.5–5% of the mode and are characterized by internal recrystallized quartz ribbons and a strong orientational fabric. They comprise phyllites, slates, and mylonitic gneisses. Unmetamorphosed and slightly metamorphosed sedimentary source rocks (Ls) are represented by lithoclasts of graywackes, quartzitic sandstones, siltstones, mudstones, and cherts. Quartz clasts with an internal, nearly equigranular and weakly recrystallized quartz grain texture were interpreted as quartzites and/or sandstones (≤2%; separately listed as QT in Table 1) and grouped together with chert fragments (C; 3–10%) as polycrystalline quartz (QP). This group is of particular interest because QT contain relics of undefined microorganisms. Opaque components are mainly represented by ilmenite, magnetite, pyrite, hematite, and organic material. The nonopaque heavy minerals consist of zircon, apatite, tourmaline, rutile, epidote, and rare pyroxene. Both accessory component groups can be found throughout the study region. The graywacke’s low degree of sorting and mean compositional maturity index of 1.25 can be interpreted as the result of intensive physical weathering and short transport distances, respectively, similar to conditions found in modern convergent margin settings. The graywackes are further characterized by a mean roundness of R = 2.69, with the majority of lithoclasts displaying subangular to subrounded shapes. In fine-gravel grain size samples, however, single grains that overstep mean grain
101
size are rounded to well rounded, which may indicate ephemeral sediment input. Depending on the degree of contact-metamorphic overprint, the mineralogical composition of the tuffaceous graywacke intercalations ranges from quartz + feldspar + epidote ± hornblende to quartz + hornblende + feldspar + diopside - ± grossular ± wollastonite (Kemnitz, 1998). They show little difference in their chemical composition (Table 2). Compared to the normal graywackes, however, an excess of Ti, Ca, and Fe minerals is present, which can be attributed to former mafic minerals typical of basic magmatic and volcaniclastic rocks. Embedded in non- to weak contact-metamorphic sequences, a sedimentary microfabric is preserved and quartz clasts, microcrossbedding, and even biotic relics may occur (Weber et al., 1990). They most probably represent diluted mass flows, derived from eroded intermediate to basic magmatic material. Structural and Metamorphic Features Cadomian folding produced large-scale, open anticlinal structures that are mainly north-vergent and striking east-west, with limb widths of some hundred meters (Schwab, 1962). Fold structures mainly seen in quarries show a clear dependence on lithology. Open folds that tend to shear along the fold axial plane were formed where coarse to medium sand beds of relatively large thickness dominate. In sections where incompetent beds dominate, however, tight parallel and similar folds accommodate up to 60% shortening (Kemnitz and Budzinski, 1994). Owing to the very low grade of metamorphism during deformation, and depending on the abundance of phyllosilicates, grain size, and primary anisotropic sedimentary features, the S1 foliation is typically disjunctive and controlled by pressure solution and rotation. Hence, a weakly developed penetrative S1 cleavage is restricted to shale beds with phyllosilicates >50% of the mode, principally sericite and chlorite. An incipient stage of S1-crenulation cleavage can be observed on a microscopic to submicroscopic scale in silty beds (Kemnitz and Budzinski, 1991). This cleavage can also be genetically related to the Cadomian folding process. The age of S1 foliation in pelitic layers of the Lausitz graywackes predates the postcollisional Cadomian granodioritic intrusions (Hirschmann, 1966). This timing can be shown by formation of a second foliation (S2) produced by contact-metamorphic overgrowth of biotite and white mica. It preferrentially occurs in laminated graywackes and siltstones with only weakly developed or practically no S1 (Kemnitz and Budzinski, 1991), where biotite tends to replace former anchimetamorphic and detrital phyllosilicates along bedding and diagenetic foliation, respectively. It is clearly not related to the Cadomian folding process and also overgrows metamorphic textures of the anatectic thermal front. In the graywackes of the eastern Lausitz region, a younger chlorite blastesis is associated with a very low-grade S-C texture visible only in thin section. This foliation, related to a more or less east–west-directed Variscan deformation stage and better
TABLE 1. LITHOCLAST COMPOSITION OF GRAYWACKES OF THE LAUSITZ GROUP (KEMNITZ AND BUDZINSKI, 1994) Region
102 Ia Ib Ib Ib Ic Ic Ic Ic Ic Id Id Id Ie I I I I I I I I I I I I I I I I II II III III III III III III III III III III III III III III III III III III III III III IV IV Dobríš Dobríš Dobríš Dobríš Dobríš Lhota Libuš
Zone 3 2 0 0 0 0 1 2 2 Encl. 2–3 2–3 2–3 2–3 2–3
Sample 821 d 7-22 7-45 978 965 966a 966b 961 962 782 c 868 a Cgl 647b 647 c
661 b
2–3
2–3 2–3 3 3 2
0–1 1–2 1(–2)
777 a 777 c
967
877 a 877 b 886 a 913 a
1–2
0–1
0–1
1–2 1–2
3 3
918 a 919 b 921 b 932 a 934 a 934 b 4755-348 4755-349 S7 S8 S12 S24 Sb22 Sb323 a 973a1 973a2 973b 973c1 973c2 975 976a
GS (µm) DetMc 460 490 375 350 600 1200 1500 500 500 590 400 420 >1 cm 425 420 390 665 650 470 480 375 425 555 620 390 435 430 390 500 350 360 530 400 315 460 380 375 330 380 380 400 355 450 400 435 465 550 775 400 435 425 440 490 480 400 1200 550 550 2000 550 500
x x x 6 48 7 7 35 30 2 2 15 0 16 x x x 11 4 x x 2 2 4 2 x x x 80 1 5 113 158 28 101 73 95 296 18 104 76 45 43 4 83 18 4 20 156 73 x 46 x x 0 40 17 76 0 x 26
QM
P
177 113 118 66 39 45 11 92 130 357 105 180 40 490 366 115 216 543 310 376 143 369 76 169 261 378 370 202 105 109 310 423 228 91 272 277 460 159 173 242 244 200 397 301 358 136 212 116 353 383 317 223 78 175 41 1 31 75 0 33 111
56 34 45 25 23 13 1 45 85 122 17 55 0 152 120 32 41 170 147 140 73 87 21 35 77 105 88 52 40 41 70 70 28 29 122 71 187 68 40 63 76 84 114 89 97 27 48 41 144 120 102 40 30 76 41 0 28 36 0 80 57
QF
Va
Vi
Vb
Kemnitz 31 42 14 23 20 24 20 4 12 19 9 8 0 5 25 27 50 30 100 55 15 38 87 27 0 96 63 66 84 39 47 17 44 62 108 102 100 27 12 55 23 27 145 91 5 27 15 36 12 54 83 46 118 51 58 38 13 14 20 19 50 45 124 101 93 59 22 36 29 71 59 72 16 79 41 65 12 48 35 64 23 96 26 41 32 89 119 45 77 51 57 41 55 27 40 32 50 68 96 79 40 108 9 34 2 12 12 16 20 14 0 0 6 10 10 16 0 1 18 15 39 28
K
13 7 8 12 31 50 28 26 30 36 116 10 46 5 31 4 53 58 10 39 6 22 16 20 30 22 12 10 18 0 28 10 22 8 29 29 13 14 14 14 8 27 8 36 16 4 8 9 21 10 29 3 1 0 37 3 66 107 28 68 41
7 2 11 44 53 38 34 68 86 12 6 5 0 5 0 4 3 23 6 54 9 38 4 14 26 11 9 5 54 0 20 7 13 8 37 37 34 12 13 7 14 14 13 21 20 3 4 9 23 17 11 5 1 1 140 14 169 221 79 248 290
5 0 1 2 0 12 0 0 0 15 3 10 0 2 1 1 4 9 6 16 4 20 8 15 38 8 6 7 0 8 22 11 4 2 11 26 6 4 6 4 7 3 1 11 19 5 17 6 16 18 1 4 0 0 24 3 0 0 0 20 180
C 15 6 11 5 24 57 57 0 0 72 11 20 1 11 28 13 19 19 14 35 7 15 11 53 23 24 16 5 8 22 14 11 13 5 26 20 4 2 62 7 7 15 23 18 31 9 26 19 11 18 22 6 0 6 15 1 0 0 0 23 35
Qt 2 0 14 8 6 8 0 2 0 98 5 0 7 0 0 0 23 27 18 0 0 0 27 0 0 0 9 1 0 2 11 0 0 0 1 4 0 0 2 0 0 0 5 0 0 0 0 0 0 2 0 0 0 0 20 4 2 1 0 0 0
M 0 0 0 0 30 6 3 3 8 5 1 2 39 0 0 0 11 15 2 7 0 3 0 1 1 1 0 1 0 0 0 6 3 3 14 10 3 6 0 10 3 6 7 1 7 0 4 7 6 6 1 5 0 0 9 2 0 0 0 0 3
S
n
0 0 6 0 5 21 22 10 14 0 3 0 107 0 4 0 0 1 0 0 0 1 4 1 0 1 0 0 5 2 0 5 1 7 1 4 2 1 2 0 0 3 0 1 6 4 6 24 9 5 0 0 0 0 60 5 55 34 33 42 103
348 199 258 186 242 267 161 298 433 872 320 396 336 794 673 233 476 1075 640 734 292 791 199 359 522 679 679 379 257 223 570 768 464 211 613 609 804 372 372 446 478 419 689 642 682 286 407 303 701 754 631 329 124 286 421 33 367 500 141 547 887
Note: New data are shaded in the table. (Ie)—pebble composition of a conglomerate enclave in granodiorite (database: Lobst, 1996); GS—mean grain size; DetMc—number of point-countered detrital micas; QM—monocrystalline quartz; P—plagioclase; K—potassium feldspar; QF—quartzfeldspar aggregate; Va—volcanic rock, acid; Vi—intermediate; Vb—basic; C—chert; Qt—quartzite-sandstone; M—metamorphic rock; S— sedimentary rock; x—detrital micas chloritized and/or recrystallized; n—number of point-counted lithoclasts/counted pebbles. Regions/sample locations: I—Western and southern surroundings of Kamenz (KM); Ia—quarry at Wetterberg, ~33 km west of KM; Ib—quarries near Königsbrück, ~16 km west of KM; Ic—quarry at Schwarzkollm, ~17 km north of KM; Id—quarry at Oßling, ~10 km northeast of KM; Ie—quarry at Kindisch, ~5 km south of KM; II—Radeberg and surroundings; III—Görlitz region (GÖ); IIIa—valley of the river Schöps, ~3–5 km west and northwest of GÖ; IIIb— eastern and southeastern surroundings of GÖ, including Poland; IV—“Gröditzer Skala,” a valley ~25 km WNW of GÖ. See also Figures 2 and 5.
The Lausitz graywackes, Saxo-Thuringia, Germany
103
TABLE 2. AVERAGES OF MAJOR, MINOR, AND TRACE ELEMENT ANALYSES OF PELITIC AND GRAYWACKE INTERVALS OF THE LAUSITZ GROUP Group
1.1
1.2
2.1
2.2
3.1
3.2
4
5
6
I
73.2 0.5 13.2 4.6 1.2 0.9 2.3 3.3 4
62.8 0.8 17.9 9.0 2.8 1.2 3.9 2.5 5
70.9 0.7 14.0 6.7 1.7 1.5 2.5 3.2 9
68.1 0.7 16.1 6.9 1.9 1.2 3.2 2.2 4
72.5 0.6 13.8 5.5 1.6 1.4 2.9 2.9 8
59.1 1.0 19.7 8.7 2.6 0.8 5.1 1.0 1
69.5 0.7 15.2 5.7 1.8 1.4 3.1 2.7 11
69.8 0.6 15.0 5.2 1.7 1.4 2.7 2.6 3
63.6 0.5 13.3 3.6 0.9 10.2 0.4 0.6 6
69 741 16 279 29 253 97 414 3977 14 2 12 81 14 15 46 29 18 10 3.4 75 2.4 31 17
75 759 15 213 31 192 124 410 4223 11 2.4 11 113 21 20 50 37 19 14 3.8 109 3 36 6
57 880 14 264 26 243 96 336 3622 6 1.8 10 73 9 14 46 27 18 10 2.9 65 2.3 31 9
105 1244 14 227 36 150 137 511 4995 20 4.3 10.5 92 62 n.d. n.d. 36 14.5 n.d. 9 104 n.d. n.d. 2
74 803 14 216 28 198 116 428 4101 17 2 10 78 22 11 35 27 16 9 4.4 81 2.2 26 12
66 726 14 232 31 207 103 452 4380 6 1.7 9 63 15 16 44 22 17 13 4 74 2.8 33 12
67 198 13 211 34 333 46 1578 5086 14 2.4 8 51 32 14 47 22 12 10 4.5 53 2.8 12 15
IIa
IIb
66.0 0.6 15.6 4.4 1.4 13.0 0.2 0.2 2
48.1 2.6 13.7 18.2 6.2 11.3 0.8 1.5 9
Major elements (wt%) SiO2 TiO2 Al2O3 FeOtot MgO CaO K2O Na2O n
64.1 0.8 17.0 8.1 2.1 1.2 3.6 2.4 6
Minor and rare earth elements (ppm) Zn Ba Nb Zr Y Sr Rb Mn Ti B Be Co Cr Cu Ga La Ni Pb Sc Sn V Yb Li n
92 796 16 204 31 141 133 445 4766 41 2.6 11 92 34 20 38 34 20 15 4 100 2.8 40 7
60 663 14 254 27 228 100 354 3544 11 2 9 64 14.5 14 40 27 17 10 3 64 2.4 30 12
87 826 12 212 32 185 144 398 4374 33 2.7 14 84 26 22 45 38 19 15 3.7 96 3.2 54 14
53 n.d. 187 271 13 5 198 251 34 86 347 120 41 9 1430 2480 3262 19,500 16 6.0 4 0.3 8 28.0 62 51.0 47 58.0 12.5 16.0 47 26.0 22 67.0 12 6.0 10 40.0 6 8.0 52 685 3 8.0 6 12 8 1
Note: Major elements are in wt%; minor and trace elements in ppm. Pelitic intervals are contact-metamorphic groups 1.1, 2.1, 3.1; and graywacke intervals are contact-metamorphic groups 1.2, 2.2, 3.2. Designations: 4, 5, and 6—Cadomian syn- to postcollisional anatexites and granodiorite intrusions; I—tuffaceous graywacke (volcanogenic mass flow members of the Lausitz Group); IIa—volcanogenic mass flow (enclaves in granodiorite and anatexite); IIb—metabasic rock (enclaves in anatexite). n—number of samples, n.d.—not determined. The numbers describe the following rock types and contact-metamorphic zones: 1.1—Zone 1: (epimetamorphic) mudstone and siltstone, non- to very weak contact metamorphism; 1.2—Zone 1: silty to fine sandy graywacke, non- to very weak contact-metamorphosed; 1.3—Zone 1: medium to coarse sandy graywacke, non- to very weak contact-metamorphosed; 2.1—Zone 2–3: metapelite (spotted shales); 2.2—Zone 2–3: metapsammite (hornfelsic graywackes); 3.1—Zone 4: metatectic to migmatized metapelites; 3.2—Zone 4: metatectic to migmatized metapsammites; 4.0—anatexites with streaky, nebulitic to homophane textures, rich in enclaves; 5.0—granodiorites and muscovite-bearing quartzdiorites with fine-grained, porphyric texture and rich in enclaves; 6.0—granodiorites; I, IIa—basic tuffaceous graywackes; IIb— metabasalts.
104
Kemnitz
known from the granodioritic rocks, can be defined as S3 (Kemnitz and Budzinski, 1991). METHODS AND DATABASE Detrital and Geochemical Composition—The Frame Each tectonic setting is characterized by its type of magmatism, sedimentation, and erosional regime. To reconstruct the plate-tectonic evolution of the study area, geochemical discrimination schemes were first applied to basic magmatic rocks and their sources (Pearce and Cann, 1973). Likewise, the tectonic setting also controls sediment composition, especially that of clastic (psammitic) sediments, through their chemical (Bhatia and Taylor, 1981; Bhatia, 1983; Bhatia and Crook, 1986; Roser and Korsch, 1986) and detrital composition (Dickinson, 1970; Dickinson and Suczek, 1979; Dickinson et al., 1982). Interpretation of classification and/or discrimination schemes becomes more reliable in sediments of lower maturity, which is reflected by a low ratio of stable to unstable components (see below). The petrographic examination of the detrital fragments is the primary method of constraining the nature of the source rocks. However, where diagenesis and metamorphism inhibit this method, whole-rock geochemical classification helps address the problem and provides additional information. As parts of the Lausitz Group experienced strong contact metamorphism, both methods have been applied and the data compared. In a first step, the progressive contact-metamorphic effects have been used to define and code the following zonation: Zone 0—no microtextural features of contact metamorphism; Zone 1— appearance of first single biotite flakes; Zone 2—biotite growth, formation of porphyroclasts in pelitic material, grain boundary changes, and subgrain growth in quartz clasts; Zone 3—start of overgrowth of internal textural features in lithoclasts, general grain growth, and onset of cordierite porphyroblastesis in pelitic material; and Zone 4—development of metamorphic granoblastic texture, incipient formation of a metamorphic foliation and mobilization of quartz layers, recrystallization of detrital feldspar clasts, and loss of internal palimpsestic textures in larger lithoclasts. Four study areas were selected (Fig. 2), all of which exhibit gradual transitions from non- or weakly contact-metamorphosed graywackes (zones 0–1) up to migmatized hornfelses (zone 4). These characterize the inner contact halo around granodioritic and anatexitic rocks: I. Kamenz region, zones 0–4; II. Radeberg region, zones 1–4; III. Görlitz/Zgorzelec region, zones 0–3, plus weak Variscan deformation; and IV. Weißenberg region, zones 2–4. Additionally, data from a number of isolated outcrops and quarries farther to the north for graywackes (in zones 0–3) have been included. A lithological section of 35 m, examined in a quarry (Oßling; Fig. 4), has been used as a reference section (Kemnitz and Budzinski, 1994).
Detrital Component Analysis The ratio of stable to unstable detrital components in siliciclastic sediments (Dickinson, 1970) is primarily controlled by the lithology of the source area and its specific resistance to erosion. Thus, climate and altitude are integrated secondary factors. Apart from lithology, such parameters as the length and duration of transport and the nature of the depositional facies are important. The resulting classification is also strongly influenced by the type of weathering (chemical versus physical) during deposition. Extreme climatic conditions and/or high altitudes can amplify these effects. In particular, the abundance of detrital micas, ore minerals, and heavy minerals is highly sensitive to these processes, such that these components are usually excluded from the analysis. To characterize the tectonic setting of a source-rock area and the basin type, immature psammitic rocks, such as graywackes and lithic sandstones, provide the best results. Eroded and transported during periods of rapid weathering and accumulation, a relative broad and representative spectrum of unstable fragments will be available. With respect to sorting effects, however, their minimum grain size should exceed 200 μm. In this article, the definitions of grain types and application of ternary systems follows those of Dickinson (1970) and Dickinson and Valloni (1980): single quartz grains (QM) and polycrystalline quartz aggregates (QP = C + QT) are the stable components, with Q = QM + QP (Table 1). Single feldspar grains (P and K) and QF form the F component (feldspar total). Rock fragments combining feldspar and/or quartz with mica and additional recognizable minerals (e.g., zircon, rutile) were either classified as QF or Ls component, depending on textural features, that distinguish an intrusive magmatic from a metamorphic source rock. Together with polycrystalline rock fragments (L) they are classed as unstable components. The Q:F:L plot (with F = P + K) is both the most commonly used and simplest ternary system for depicting the stable-:unstable-component ratio. However, it lacks important information on source rock lithology and genesis because those Q with QP components have completely different origins. QM and F represent the magmatic component in the form of volcanic and intrusive source rocks, respectively. The type of magmatic source rock can be further defined with the help of the P:K ratio in the QM:P:K diagram. Furthermore, the presence of feldspars of different composition in association with either quartz grains of either QM or QP type, or heavy minerals of possible metamorphic origin, and/or detrital micas (QF components) can help to determine whether granitoids or basement rocks played a larger role. L is subdivided into volcanic fragments (Lv) and Ls. Lt then comprises the sum of volcanic, sedimentary, and metamorphic rock fragments, which includes QP and QF. Application of the QM:F:Lt diagram therefore allows for a more detailed genetic interpretation than the Q:F:L diagram. In our case, the QP:Ls:Li diagram provides further discimination, with Li representing all igneous rock fragments, including QM and F. Point counting of detrital compositions was performed on fifty-four thin sections from Lausitz graywackes (Table 1) and
m
188/57
b
b
c
a
a
a
a
e
d
c
m
m
19
20
21
22
25
Mineralized veins and lodes
195/55 190/50
190/45 195/55
195/55
190/45
190/70 195/55
strongly jointed
Inversion of stratigraphic facing (hinge areas of folds)
Direction and angle of bedding and fault plane dipping
Intraclasts
200/60 215/50
205/40 200/45 200/45
S fs mscs
2m debris 180/70 0.5 m
M
Fault planes 167/58
d
c
b
d
c
b
d c d b c d
c
d b
b
c
a e d
Turbiditic intervals (Bouma, 1962) Favososphaera conglobata sensu Burmann (Burmann,1972 a)
200/47
200/55
195/45
192/55
S fs mscs
a-e
M
Intercalation of a volcanogenic greywacke
26
27
28
29
30
Load marks
Cross stratification
Planar bedding and lamination
55/32
167/58
188/47
173/80
177/82
(Hg. wall)
190/57
S fs mscs
Legend:
M
d b
b e
d
c
b
d
c
?
b e b
d
c e b c
e
d
c
m
13
14
15
16
17
18
C
180/50
185/50
200/35
180/50 180/50
190/45
185/43
S fs mscs
d
c
c d
d
c
d b d
b
c
8
9
10
11
12
m C
180/55
180/45
185/55
190/40
185/47
170/50
190/47
185/55
S fs mscs
b
b
d
c
b
c
d
d c
c
d
C
185/60
185/60
190/60
185/65
185/60
185/55
180/65
180/60
180/57
190/50
180/45
185/50
185/40
190/50
S fs mscs
(Hanging wall)
0
1
2
3
4
5
6
7
m
b
d
c
d
c
b
d
a
d
c?
e
d
c
b
c
b c d
c
d b
Figure 4. Lithological section of the Lausitz Group, region Id (Kamenz area, quarry Oßling) beginning down right. Note inversions of stratigraphic facing caused by Cadomian folding. Postsedimentary mineralization resulted from a nearby granodiorite intrusion. Grain-size scale of the graywackes (head): M—mudstone; S—siltstone; fs—fine sand; ms—medium sand; cs—coarse sandstone. Hg. wall—hanging wall.
31
32
33
34
35
The Lausitz graywackes, Saxo-Thuringia, Germany 105
106
Kemnitz
from three other Saxo-Thuringian basement samples (Leipzig Group, Katzhütte Group; Table 3). Additionally, seven thin sections from the Teplá-Barrandian Neoproterozoic (Šteˇchovice Group) were point-counted (Table 1). On average, between two hundred and five hundred points (lithoclasts and/or minerals) were counted per thin section. Larger numbers of points were counted for samples with a finer grain size or higher degree of contact-metamorphic overprinting. The point-counted fragments and grains were classified into the categories listed above using the traditional method defined by Ingersoll et al. (1984). To maximize information, each grain within the range of the mean grain size in a thin section was counted, without maintaining a strict regular counting grid. The grain sizes of the examined graywackes range from medium to
coarse sand (350–1000 μm) to occasional fine gravel (2–6 mm). Hence, the data are statistically comparable. Because the number of less-stable components increases with increasing grain size, including the larger grain-size classes allows a more complete diagnosis of the source area. This effect occurs above the boundary of coarse-sand to fine-gravel grain size. Sorting measures and roundness classes for each type of lithoclast were also recorded, and grain-size intervals were selected based on the mean grain size in each thin section, within pre-established interval limits. The degree of roundness (R) was evaluated visually and grouped into classes I–V (angular, subangular, subrounded, rounded, and well rounded). R is primarily controlled by either transport processes or lithology. Secondary factors, such as contact-metamorphic overprint, however,
TABLE 3. DETRITAL COMPOSITION OF NORTHWEST SAXONIAN (LEIPZIG GROUP) AND THURINGIAN (KATZHÜTTE GROUP) NEOPROTEROZOIC GRAYWACKES Region Jena Jena Jena Jena Jena Jena Jena Jena Jena Gump Leut Zeitz Zeitz Zeitz Zeitz Zeitz Pegau Pegau Pegau Groitzsch Störmthal Störmthal Kitzen Leipz-Delitzsch Leipz-Delitzsch Leipz-Delitzsch Leipz-Delitzsch Leipz-Delitzsch Leipz-Delitzsch Leipz-Delitzsch Leipz-Delitzsch Leipzig Leipzig Katzhütte Katzhütte Katzhütte
Sample
977 SB512a SB512b
QM
P
K
QF
Va-i
55 54 54 34 83 48 46 45 48 48 51 53 55 53 50 52 53 50 50 53 59 41 43 52 50 54 46 50 58 52 55 61 48 57 36 40
12 10 12 6 5 14 7 13 12 9 10 12 11 9 10 10 12 10 13 12 12 13 16 10 14 9 12 9 15 13 10 20 8 11 40 32
18 18 16 14 10 18 18 13 16 25 19 17 17 18 15 14 13 19 17 14 16 24 22 18 4 18 15 17 8 18 14 7 17 17 2 0
3 3 5 3 1 4 7 6 4 3 4 6 2 5 6 4 3 5 5 4 3 10 7 6 8 8 5 4 4 4 11 8 4 4 6 10
1 1 1 2 0 3 5 4 2 2 3 2 2 3 2 3 2 2 3 4 1 2 2 2 3 0 4 3 6 1 1 14 1 1 5 6
Vi
42
6 4
Vb
S
C
Qt
M
1 1 1 1 0 2 1 2 2 1 1 1 1 1 1 1 1 0 2 2 1 1 1 1 1 0 3 2 1 1 0 3 1 0 1 1
4 4 3 27 0 3 6 5 6 6 5 3 5 5 4 5 6 4 4 2 1 2 3 5 3 0 6 6 1 5 1 0 7 1 0 1
3 4 3 2 0 3 4 6 4 1 2 1 3 2 4 4 3 3 3 2 2 3 2 2 4 6 4 2 1 2 1 10 3 0 2 3
2 4 3 8 1 4 5 3 3 2 3 3 2 3 5 4 4 3 2 4 4 4 3 3 11 5 3 3 6 2 7 0 7 7 2 3
1 1 2 3 0 1 1 3 3 3 2 2 2 1 3 3 3 4 1 3 1 0 1 1 2 0 2 4 0 2 0 0 4 2 0 0
Note: Calculated on the basis of 100%. Data from Sehm (1972) and Kemnitz et al. (1999), including new data (shaded). QM—monocrystalline quartz; P—plagioclase; K—potassium feldspar; QF—quartz-feldspar aggregate; Va-i—volcanic rock, acid to intermediate; Vb—basic; Vi—intermediate; S—sedimentary rock; C—chert; Qt—quartzite-sandstone; M—metamorphic rock.
The Lausitz graywackes, Saxo-Thuringia, Germany may lead to major grain boundary changes. The graywackes were therefore grouped according to their contact-metamorphic zones, each being represented by five to nine samples (Table 1). With respect to recrystallization effects, provenance interpretation has been based on graywackes from zones 0–2/3. Geochemical Whole-Rock Classification Geochemical analysis of the sediments was performed on two separate groups of samples: (1) the psammitic parts of graywacke from intervals of fine-sand to fine-gravel, and (2) pelitic material from intervals of mudstones to siltstone showing all degrees of contact metamorphism. In this way, the effects of metamorphic overprint could be tracked and the resulting changes in the chemical composition distinguished from other causes. Using the scheme of Roser and Korsch (1986) allows psammitic rocks of differing degrees of maturity to be distinguished: (1) increasing SiO2 content mirrors increasing maturity (by sorting processes) and correlates with decreasing K2O and Na2O; (2) high alkali content reflects derivation from feldspars of acid to intermediate magmatic source rocks, including orthogneisses; (3) Al2O3 correlates negatively with the sediment maturity (Bhatia and Crook, 1986); (4) high content of mafic major elements (Fe2O3, MgO, CaO, and TiO2) indicates basic magmatic source typical of rift or volcanic arc settings; and (5) minor and rare earth elements (REE; e.g., Ti, Zr, V, Sc, La, Y; Bhatia and Crook, 1986; McLennan, 2001), which behave stably, best define the tectonic setting of the source area and basin (Tables 2, 4, and 5). Mineral Chemistry as a Tool for Source-Rock Classification The source-rock classification of detrital micas (Wybrecht et al., 1985; Brigatti and Gregnanin, 1987) uses the relationship between formation temperature and the chemistry of the mineral. With changing environmental P/T conditions, cation exchange takes place between the tetrahedral (Si, AlIV) and octahedral (Mg, Fe2+, AlVI) sites (Velde, 1967; Guidotti and Sassi, 1976, 2002). But the intensity of deformation (i.e., the degree of recrystallization) also plays an important role (Behrmann, 1984). In sheared, but only slightly recrystallized domains, the micas may still preserve their primary temperature memory. In this case, grain size is important because larger grains are less susceptible to reaction. Under low-grade conditions, the first metamorphic growth of white micas takes place along cleavage planes before the larger detrital micas start to recrystallize (Stephens et al., 1979). The larger detrital micas consequently tend to preserve their primary chemistry during metamorphic temperatures above 350 °C (c.f. Villa, 1998). Detrital biotite and white mica grains from the Lausitz Group were analyzed by microprobe (Tables 6 and 7). Samples range from noncontact metamorphosed graywackes (below chlorite-in) to graywacke hornfelses above the cordierite-in boundary in associated metapelites (Kemnitz et al., 1994). In all cases, the detrital micas were easily distinguished by their
107
characteristically deformed, bended shapes and grain size (350–450 μm on average). Analysis was performed in polished and carbon-coated thin sections with a CAMECA-SX50 microprobe operating in wavelength-dispersive mode (micas, feldspars in part; Table 8) and with a scanning electron microscope (SEM) (Zeiss, DSM 962) equipped with an energy-dispersive system (feldspars) and a NORVAR Si-Li detector (Thermo Electrons). The beam current of the microprobe was 20 nA, and the acceleration voltage was 15 kV. Matrix correction was made with a PAP program. The counting time was 20 s on average using a beam diameter between 2 and 8 μm, depending on the analyzed mineral. For the SEM, the beam current was ~0.33 nA at 20 kV acceleration voltage. Calculations were performed using ASTIMAX and MAC mineral references, and a PROZA/ZAF correction program. RESULTS AND INTERPRETATION Provenance and Tectonic Setting of the Lausitz Group Graywackes The mean composition of graywackes in each study region is depicted in the form of pie charts (Fig. 5). The only significant variation these show is due to two factors: First, pie charts Ic and Ie, with their high percentage of unstable lithoclasts, represent a single gravel-size and conglomerate sample, respectively. Both samples consequently exceed the coarse-sand boundary (>1 mm), which leads to a large increase of unstable components (mainly Ls and Lm) relative to stable components (QM, QP, and the F component). In the discrimination diagrams (Fig. 6), the conglomerate sample is therefore marked by a large circle, to avoid misinterpretation of this grain-size controlled shift. Second, the detrital composition reflects the effect of secondary reduction of unstable lithoclasts caused by contact-metamorphism. In Figure 5, this relationship is most clearly seen in samples from region IV, which experienced the strongest metamorphic overprint, and those from region II, which also experienced a higher grade of contact metamorphism. In Figure 6A, P is conservatively calculated as 100% of F (with F = P + K). Following the definition (see above), all plagioclase appears as an intrusive source-rock component. However, if ~50% of P was added to the volcanic-rock component (intermediate and basic), which would be a more realistic assessment, the discrimination fields of the systems QM:F:Lt and Lm:Lv: Ls would shift closer to the volcanic and magmatic arc areas. Applied to the Lm:Lv:Ls diagram, which is used to discriminate between basin settings (Ingersoll and Suczek, 1979), this conservative calculation produces a “mixed” depositional setting between fore-arc and back-arc basin. A shift to the magmatic-arc field, however, would show a fore-arc basin setting. The magmatic arc influence is better reflected in the QM:F: Lt and QP:Li:Ls diagrams plotted in Figure 6B, where plutonic rock derivation is strained and a dissected arc provenance with a large amount of continental input (QM-F-Lt) is shown.
66.9 16.8 0.75 1.23 4.1 1.2 1.92 3.19 3.3 0.18 0.4 1
66.5 16.8 0.77 1.96 3.9 2.4 1.65 3.78 2.1 0.19 0.5 1.8
Zn Ba Nb Zr Y Sr Rb Mn Ti B Be Co Cr Cu Ga La Ni Pb Sc Sn V Yb Li
81 977 15 217 30 247 127 513 4500 19 2 12 50 14 17 32 27 14 11 4 62 2 46
94 682 16 199 30 197 151 510 4540 45.3 3.1 13.6 77.5 29.2 20.9 30.3 37.3 20 16 4.3 93.3 3 30
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
750 I
92 849 17 204 35 152 150 521 4740 53 2.7 11.3 139.4 31.5 27.7 46 43 20.3 18.9 4.6 126.7 3.3 48
Minor and rare earth elements (ppm)
SiO2 Al2O3 TiO2 Fe2O3 FeO MgO CaO K2O Na2O P2O5 H2O+ H2O–
744 I
109 776 17 186 32 97 161 572 5210 1.8 1.2 9.5 100.6 9.7 12.5 41.1 29.9 15.5 10.2 2.2 76.2 2.2 60
61 19.5 0.88 1.52 5.6 2.8 1.15 4.41 1.3 0.15 0.6 2.7
768 I
47 876 15 212 22 125 121 366 5030 78 3 16 90 60 21 38 36 18 13.4 5 132 2.7 32
63.4 17.3 0.81 5.44* 4.86 1.8 0.49 3.79 2.5 0.13 n.d. n.d.
867 I
162 727 14 184 37 95 100 415 5180 38 2.6 9 95 70 23 42 37 34 18.2 3.8 122 3.4 44
61.2 16.5 0.79 6.89* 6.3 2.6 1.4 2.95 2.2 0.4 n.d. n.d.
871 I
62 687 15 229 29 75 118 221 4160 50 3.3 8 88 26 n.d. n.d. 31 19 n.d. 4.1 88 n.d. 20
65.6 15.2 0.71 5.19* 4.67 1.9 0.32 3.25 2.7 0.15 n.d. n.d.
S 24 III
62 458 21 465 38 247 75 439 4990 13.8 1.8 9.5 144.2 11.1 14.9 82.8 24.4 15.5 12 3.4 91.8 3.1 37
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
1.2 647c I
69 782 15 279 31 277 101 357 3850 10.3 2.2 10.7 94 17.7 17.2 51.4 30.9 16.1 11 3.4 77.8 3 33
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
684 I
50 800 16 266 26 239 106 388 3610 10.7 2.6 9.1 50.6 10.7 17.6 41.7 26.2 18.9 9.6 3.9 55.8 2.5 32
72.2 14 0.61 1.09 3 1.4 1.38 3.03 3 0.14 0.4 0.7
742 I
70 685 13 212 27 269 94 426 3690 8.7 2.4 11.3 50.8 6.6 15.9 35.4 28.9 18.8 10.1 3.4 62.8 2.5 35
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
743 I
75 628 14 191 24 215 109 405 3650 19.3 3 11.6 63.8 19.5 20.6 35.1 34.5 21.7 11.3 3.9 72.6 2.4 34
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
746 I
41 640 15 274 27 232 77 359 3580 10.6 2 8.7 57.3 12.9 15.2 40.8 27 17.1 10.6 2.9 70.4 2.4 29
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
748 I
TABLE 4. MAJOR AND TRACE ELEMENT ANALYSES OF PELITIC INTERVALS AND GRAYWACKE INTERVALS OF THE LAUSITZ GROUP
1.1 741 I
Major elements (wt%)
Group Sample Region
Continued
56 676 14 254 26 235 93 367 3550 10.6 1.4 8 53.9 12 11.9 39.5 26.8 13.6 9.4 2.4 56.9 2.3 32
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
749 I
108 Kemnitz
Zn Ba Nb Zr Y Sr Rb Mn Ti B Be Co Cr Cu Ga La Ni Pb Sc Sn V Yb Li
(ppm)
SiO2 Al2O3 TiO2 Fe2O3 FeO MgO CaO K2O Na2O P2O5 H2O+ H2O–
(wt%)
Group Sample Region
80 373 13 219 25 203 103 409 3460 12.7 2.2 10.7 47.6 15.7 12.8 33.6 30.5 17.1 8.7 2.3 50.5 2 35
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
751 I
48 961 14 179 26 269 234 348 3260 1.9 1.3 7 72.9 27.9 19.1 24.9 20.5 18 10 3.7 82.7 2.3 23
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
755 I
59 626 12 174 23 238 79 363 3050 8.9 2.3 10.8 47.9 13.6 15.5 30.6 28.8 18.6 7.9 2.8 49.8 2 33
73.2 14.1 0.52 1.03 2.6 1.6 1.21 2.08 3.6 0.14 0.2 0.9
769 I
50 667 14 319 24 153 66 84 2590 5.7 1.4 4.6 37.4 12.7 11.1 26.7 23.4 14.6 5.6 2.3 46.7 2.1 22
76.5 12.2 0.44 1 1.5 0.3 0.5 1.87 3 0.13 n.d. n.d.
778 I
61 664 12 211 26 159 66 306 3250 22 1.8 6 49 13 n.d. 33.5 20 18 8.8 2.6 50 2.1 17
70.9 12.3 0.54 3.86* 3.47 1.4 0.66 2.26 3.6 0.11 n.d. n.d.
878 III
54 697 16 238 25 256 90 370 3500 4.2 1.7 10.1 54.2 4.9 11.9 45.1 28.1 16.9 9.2 2.6 57.8 2.4 20
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
2.1 640 I
60 753 17 264 29 290 92 353 3750 6 2 10 80 12 14 43 28 19 10 2 65 3 24
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
659 I
92 942 16 205 40 166 200 437 5230 9 3 14 79 39 32 58 41 25 20 4 90 4 69
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
662 I
62 751 15 213 32 189 132 504 5320 4.7 3 11.1 109.8 27.6 26.1 47 34.9 16.9 17.6 4.5 126.7 3.2 52
63.1 18.8 0.86 1.57 4.5 2.5 1.3 3.49 2.4 0.14 n.d. n.d.
711 I
41 692 13 175 33 152 144 184 4170 51 4 11 105 30 27 44 37 18 16 4 125 3 45
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
716 I
108 727 14 184 35 167 142 242 3640 49.1 2.6 11.5 91.8 28.2 24.4 36.1 48.9 20.1 12.3 3.9 129.8 2.9 52
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
718 I
101 962 16 182 29 144 160 542 4710 78 3 12 105 32 22 50 38 20 22 5 140 4 72
62.7 19.5 0.82 2.15 4.5 2.5 0.91 4.01 2.2 0.15 0.7 2.4
747 I
89 962 16 225 29 132 157 347 4390 17.2 2.3 12.8 92.6 16.2 21 42 35.6 18.4 12.8 3.8 91.7 2.7 60
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
753 I
TABLE 4. MAJOR AND TRACE ELEMENT ANALYSES OF PELITIC INTERVALS AND GRAYWACKE INTERVALS OF THE LAUSITZ GROUP (continued)
Continued
129 830 16 195 35 167 175 464 4720 46 3 15 96 17 20 46 51 19 18 4 96 3 100
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
815 I
The Lausitz graywackes, Saxo-Thuringia, Germany 109
Zn Ba Nb Zr Y Sr Rb Mn Ti B Be Co Cr Cu Ga La Ni Pb Sc Sn V Yb Li
(ppm)
SiO2 Al2O3 TiO2 Fe2O3 FeO MgO CaO K2O Na2O P2O5 H2O+ H2O–
(wt%)
Group Sample Region
102 1209 n.d. 280 28 215 172 528 4062 31 3.2 17.8 52 19.7 24 45.8 29 19 16.9 n.d. 78.9 n.d. 64
62.9 16.3 0.9 5.80* 5.22 3.2 1.4 4.1 3.4 n.d. n.d. n.d.
951 I
100 1241 n.d. 220 28 198 181 547 4062 54 3.4 19.2 48 39.1 24 45.9 34.7 17 14.5 n.d. 89.3 n.d. 71
62.2 17.3 0.7 6.50* 5.85 3 1.1 4.5 2.4 n.d. n.d. n.d.
952 I
109 550 n.d. 170 33 143 120 352 4800 87 3 21.7 68 45.3 23 43 41.9 26 16.6 n.d. 108.8 4.2 46
63.1 17.4 0.8 6.00* 5.82 2.7 1.1 3.5 2.3 n.d. n.d. n.d.
954 I
83 491 15 217 31 216 105 407 4270 6 2 13 80 36 19 40 40 22 11 3 38 3 30
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
667 II
86 681 17 203 34 159 151 290 4610 22 2 12 115 24 18 50 40 16 17 4 110 3 53
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
793 II
114 724 18 180 34 169 186 528 5150 42 3 14 230 31 29 46 50 22 22 6 175 4 46
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
2.2 639 I
66 695 20 480 35 264 108 425 4800 9.9 2.2 9.5 168.1 10.7 17.4 67.6 26.9 16.9 11.4 3.8 87.1 3.1 33
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
649 I
76 568 17 278 35 364 96 625 5570 1.9 1.6 15.4 51.4 4.5 11.9 48 31.2 16 13.1 2.6 73.9 3.2 37
67.6 15.4 0.93 1.25 4.2 2.1 2.7 2.21 3.3 0.17 0.1 1.1
710 I
73 775 16 284 30 278 88 434 4520 13.9 1.7 11.7 96.7 11.1 15.5 42.4 32.2 13.9 1.1 3.1 92.3 2.8 39
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
792 I
53 666 14 290 26 259 76 466 3640 1.5 1.8 8.8 53.1 7.6 9.7 58.2 25.5 14.6 9.8 2.8 54.7 2.4 34
73.7 13.2 0.62 1.48 2.6 1.5 1.73 2.37 3.5 0.14 0.3 1
824 I
87 918 n.d. 280 25 269 98 479 4016 25 2.4 15.3 55 19.7 20 47.5 31.3 22 12.6 n.d. 85.1 n.b 48
66.8 15.6 0.8 5.10* 4.59 1.8 2 2.9 3.3 n.d. n.d. n.d.
953 I
81 632 n.d. 180 14 227 68 260 2561 13 1.8 8.1 40 16.2 12 24.9 10 20 7.6 n.d. 64.2 n.d. 26
74.9 12.6 0.5 3.20* 2.88 1.2 0.8 1.8 3.5 n.d. n.d. n.d.
955 I
55 739 14 249 24 266 94 510 3320 1.6 1.6 26.4 57.9 17.9 11.7 38.9 28.3 14.1 9 2.4 65.5 2.3 27
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
644 II
TABLE 4. MAJOR AND TRACE ELEMENT ANALYSES OF PELITIC INTERVALS AND GRAYWACKE INTERVALS OF THE LAUSITZ GROUP (continued)
Continued
59 514 15 293 31 306 84 395 3660 2.8 1.5 11.3 71.6 7.2 12.9 54.5 28.6 17.8 10.4 2.8 68.6 2.7 13
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
695 II
110 Kemnitz
Zn Ba Nb Zr Y Sr Rb Mn Ti B Be Co Cr Cu Ga La Ni Pb Sc Sn V Yb Li
(ppm)
SiO2 Al2O3 TiO2 Fe2O3 FeO MgO CaO K2O Na2O P2O5 H2O+ H2O–
(wt%)
Group Sample Region
63 769 12 246 29 260 86 327 3380 10.6 2 10.5 91.8 10.7 15.9 50 30 22 9.2 3.4 69.2 2.6 24
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
698 II
61 826 15 313 31 252 93 433 4030 12.1 1.7 10.6 67.6 10.3 14.5 48.9 27.4 14.6 9.3 2.6 66.3 2.6 32
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
759 II
66 863 16 240 28 228 112 346 3820 4.6 2.1 9.6 65.3 11.8 16.1 41.5 31.8 14.6 9.8 3.4 67.6 2.6 28
71.6 14.7 0.65 1.95 3 1.7 1.21 2.79 2.7 0.13 0.4 1.1
763 II
59 526 12 197 24 211 79 321 3030 4.7 1.6 11 66.7 15.1 13.3 35.4 28.8 17.8 8.7 2.9 53.6 2.3 33
74.7 13.3 0.52 1.24 2.5 1.6 1.31 2.18 2.8 0.12 0.1 1.1
764 II
51 828 15 288 29 242 91 323 3530 8.7 2.1 10 57.2 10.3 12.6 43.1 31.8 16.4 9.2 3.1 64.4 2.7 22
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
794 II
68 830 17 349 31 270 75 435 4270 9 2 12 59 9 9 50 26 16 14 4 50 2 21
71.9 13.9 0.71 2.07 2.6 1.7 1.59 2.36 3 0.13 0.5 1.1
798 II
79 782 16 226 30 204 117 352 4050 39 3.9 13 77 28 n.d. n.d. 31 30 n.d. 4.8 73 n.d. 38
67.1 14.5 0.67 5.07* 4.56 1.7 1.03 3.02 3.4 0.15 n.d. n.d.
S 11 III
65 937 16 366 38 237 98 381 4270 30 3 8 69 15 n.d. n.d. 21 22 n.d. 3.3 65 n.d. 32
69.4 12.9 0.7 4.93* 4.44 1.6 1.01 2.85 3.2 0.14 n.d. n.d.
S 12 III
72 613 15 219 32 235 109 385 3800 23 1.8 12.2 60.4 21.1 19.1 38.9 27.5 20.9 11.2 3.4 80.6 2.7 26
69 15.6 0.66 1.31 3.4 1.8 1.59 2.81 2.7 0.14 0.3 1.5
3.1 714 I
49 644 15 180 27 131 114 237 3990 14.2 2.6 9.9 85.3 34.7 20 37.7 30.4 10 13.3 4.1 141.5 2.5 28
65.2 17.6 0.73 2.54 3.1 2 0.65 3.47 1.9 0.1 0.5 2.7
715 I
103 763 13 197 36 170 194 523 4570 8.6 3.4 13.5 141.9 23.1 30.5 58 49.3 30 18.6 5 128.1 4 65
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
692 II
70 895 14 222 32 208 109 431 4660 10 2 12 124 26 14 50 40 14 15 4 107 3 28
66.9 16 0.76 1.91 3.8 2.1 1.28 3.13 2.1 0.17 0.6 1.9
795 II
82 781 18 226 35 200 109 457 4420 7.5 2.8 12 182.1 18.1 21.3 73 43.2 24.4 16.5 3.7 122.2 3.3 30
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
702 II
TABLE 4. MAJOR AND TRACE ELEMENT ANALYSES OF PELITIC INTERVALS AND GRAYWACKE INTERVALS OF THE LAUSITZ GROUP (continued)
Continued
76 856 14 235 25 205 111 424 3900 1.8 1.6 7.8 85.5 4.9 13.6 43.7 31.2 16.6 10.4 2.5 74.4 2.6 42
71.2 15.2 0.68 1.57 2.8 1.7 1.21 3.19 2.2 0.11 0.5 1.5
758 II
The Lausitz graywackes, Saxo-Thuringia, Germany 111
41 920 12 253 27 236 93 345 2820 25.8 3.2 8.4 31.2 5.6 11.2 38.5 17.8 18.5 6.3 3.3 42.2 2.3 28
74.7 13.3 0.5 1.2 2.3 1.4 1.34 3.11 2.8 0.12 0.2 0.6
3.2 709 I
49 805 15 219 24 199 103 313 3140 3.9 1.9 8.7 34.4 3.3 12.1 36 23.1 19.5 7.3 2.6 46.9 2.4 28
72.7 14.1 0.54 1.27 2.6 1.6 1.18 3.05 2.7 0.1 0.2 1
720 I
61 816 16 335 34 259 98 298 3770 6.4 2.1 10.8 111.7 7.2 19.3 61.2 26.9 22.6 10.8 3.8 76.1 2.9 n.d.
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
670 II
46 772 13 249 24 243 82 237 3250 1.5 0.7 7.5 59.1 12.1 12.2 44.3 15 17.7 7.8 2.7 65.5 2.3 28
75 12.5 0.56 1.31 2.3 1.3 1.11 2.49 3 0.08 0.2 0.7
642 II
66 940 16 313 28 232 119 314 3940 3.6 2.1 11.4 91.8 2.8 14.9 48.1 30.5 17.6 10.1 3 74 2.3 38
71.9 13.9 0.67 1.22 2.8 1.6 1.45 3.04 2.6 0.11 0.4 1
757 II
62 1190 14 302 24 243 93 345 3870 1.8 1.2 9.5 100.6 9.7 12.5 41.1 29.9 15.5 10.2 2.2 76.2 2.2 30
72.2 14 0.67 1.41 2.7 1.5 1.38 2.49 3 0.11 0.4 1
760 II
60 1180 14 275 28 271 101 437 4090 4 2 11 96 4 14 42 30 16 11 3 74 2 33
70.8 14.4 0.66 1.02 3.4 1.5 1.41 2.96 3.6 0.1 0.3 0.6
761 II
74 489 11 218 26 303 91 400 3800 1.2 0.8 11.4 55.9 32.5 7.8 61.5 34.6 13.5 11.3 2.3 52.6 2 28
71.3 14.5 0.6 1.13 3.3 2.1 2.27 1.97 2.9 0.12 0.3 0.8
801 II
51 808 11 216 22 203 82 332 3920 4.3 2.5 11.7 72.8 6.6 17.7 40.6 35.6 23.7 10.8 2.8 81.1 2.1 33
71.4 13.7 0.65 2.23 2.5 1.6 1.11 3.7 2.3 0.11 0.6 1.5
789 II
73 788 11 226 29 200 118 410 3810 19 3.6 6 69 10 n.d. n.d. 17 21 n.d. 8 66 n.d. n.d.
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
842 So
137 1700 16 228 43 100 155 612 6180 20 5 15 115 115 n.d. n.d. 54 8 n.d. 10 143 n.d. n.d.
59.1 19.7 0.96 8.20* 5.7 2.6 0.84 5.09 1 0.24 0.6 3.4
4 843 So
Note: Major elements are in wt.% and trace elements in ppm. Pelitic intervals are contact-metamorphic groups 1.1, 2.1, 3.1; graywacke intervals are contact-metamorphic groups 1.2, 2.2, 3.2. SO—biotite-hornfels raft within anatexite (location Sora). *FeO total.
Zn Ba Nb Zr Y Sr Rb Mn Ti B Be Co Cr Cu Ga La Ni Pb Sc Sn V Yb Li
(ppm)
SiO2 Al2O3 TiO2 Fe2O3 FeO MgO CaO K2O Na2O P2O5 H2O+ H2O–
(wt%)
Group Sample Region
TABLE 4. MAJOR AND TRACE ELEMENT ANALYSES OF PELITIC INTERVALS AND GRAYWACKE INTERVALS OF THE LAUSITZ GROUP (continued)
112 Kemnitz
The Lausitz graywackes, Saxo-Thuringia, Germany TABLE 5. RARE EARTH ELEMENT ANALYSES FROM METAPELITIC GRAYWACKE INTERVALS OF THE LAUSITZ GROUP Group Sample Region
1.1 867 Id
1.1 871 Id
2.1 659 I
2.1 954 I
2.1 667 II
3.1 692 II
5.93 38.00 5.37 69.60 8.30 31.60 6.06 1.28 4.83 0.80 4.15 0.89 2.73 0.38 2.70 0.45 9.67 2.72
3.65 42.00 4.84 81.00 9.56 36.50 7.37 1.63 6.25 1.03 5.98 1.20 3.47 0.49 3.40 0.52 10.80 3.37
4.34 43.00 6.05 73.10 8.36 31.10 5.75 1.19 4.79 0.75 4.15 0.82 2.48 0.35 3.00 0.39 12.30 2.66
1.31 43.00 4.26 66.90 7.65 29.10 6.01 1.15 5.57 0.95 5.93 1.34 4.10 0.61 4.20 0.63 7.36 1.83
6.33 40.00 5.32 80.00 9.43 35.40 6.91 1.56 5.97 0.93 5.25 1.07 3.13 0.43 3.00 0.46 12.90 3.29
11.40 58.00 5.08 96.30 11.40 43.30 8.27 1.78 7.14 1.14 6.70 1.31 3.65 0.55 4.00 0.58 15.00 3.46
(ppm) Cs La Hf Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Th U
Note: Numbers and abbreviations as for Table 8.
The REE pattern of the Lausitz graywackes (c.f. Linnemann and Romer, 2002; Table 5, Figure 7) matches other late Proterozoic to Cambrian graywackes of continental arc position (reference P40136, c.f. Taylor and McLennan, 1985). In pelitic samples, the REE abundances plot between the averages for postArchean Australian (PAAS reference) and European shales (ES reference) (Taylor and McLennan, 1985). The Na2O value for the Lausitz graywackes corresponds well to their P:K and Na2O: K2O ratios ≥1. The latter ratio is also typical of modern sands in convergent continental-margin settings (Roser and Korsch, 1986; Fig. 8B) and can be linked to older basement input. Compared to average active-margin sandstones and mudstones (McLennan, 2001) increased values of Cr, Ni, Zr, Ba, and La in both the pelitic and psammitic group play a role to some degree. V and Co show slightly lower values, which would seem to exclude the existence of larger basic arc volcanic or oceanic crust complexes. Average Cr and Ni values, however, are typical of granitic rocks of high Ca content, as is the case for Ba, Zr, and La. This kind of source rock is thought to reflect a pre-Cadomian provenance, the existence of which has been demonstrated from zircon dating by Tichomirowa et al. (2001) from several SaxoThuringian regions, including the Lausitz. However, enriched Ba may also be related to a marine environment, and Zr is generally enriched in psammitic rocks (McLennan, 2001). So although detrital composition and geochemistry demonstrate a dissectedarc source, derivation from pre-Cadomian granitic rocks remains speculative. Comparing element values from different graywacke members from zones 1–3 (Tables 2 and 4; Fig. 9), their element behav-
113
ior is lithologically controlled and specific for each of their mineralogical compositions up to low-grade contact metamorphism. This behavior is imaged in a zigzag curve (Fig. 9) distinguishing between pelitic (1.1, 2.1) and psammitic graywackes (1.2, 2.2). With increasing contact metamorphism, which becomes an important second factor in zone 3, the chemical composition of graywackes (3.1, 3.2) tends to homogeneization. Granitoids/Anatexites The anatectic thermal front that also affected parts of the Neoproterozoic basin sediments is evident in transition zones. Here, contact-migmatization can be traced from graywacke beds with different degrees of thermal overprint into the inner contact of adjoining anatectic rocks or intruded granodioritic bodies. The geochemical element distribution pattern of the homogeneous type of anatexites is strikingly similar to that of their sedimentary roof (group 6 in Table 2; Fig. 9). Hence, the anatectic melts can have been only partly mobilized from the enveloping Lausitz Group. Addition of pre-Cadomian basement to the Cadomian melt is indicated by the distribution of Nd and Sr isotopes (Linnemann and Romer, 2002), which also underline the role of Cadomian basin sediments. Along the northern to northeastern margin of the Lausitz region, the granodioritic bodies often show discordant, relatively narrow intrusional contacts with the graywacke source rock. Toward the central and southern anatectic domain, however, contact relationships appear obscured. In places where distinct intrusional contacts exist between the different types of anatexites and intruded granodioritic bodies, a sequence of multiple melt formation and intrusion can be determined: (1) Anatexites (group 4 in Tables 2 and 9), (2) enclave-rich granodioritic to quartzdioritic rocks (group 5), and (3) granodiorites (group 5). Reflecting an increasing degree of melting and temperature, this chronological sequence is enclosed in distinct zircon morphologies (Tichomirowa et al., 2001). Indications of the relatively high level of crustal emplacement of the granodiorites and anatexites are their per-aluminous chemistry, high-temperature (HT) mineral associations, numerous bedrock enclaves (Eidam, 1988; see also Table 2 and Fig. 9), and large numbers of inherited zircons (Tichomirowa et al., 2001). Metabasic Rocks Of the two principal basic rock types I and II, type IIb forms boulder-like enclaves and lenses in various migmatized and intrusive granodioritic rocks and represents fragments of volcanic arc and oceanic crust. In the Y-Nb-Zr discrimination diagram (Fig. 10C), type IIb falls within the normal mid-oceanic ridge basalts (MORB) and volcanic arc basalts (c.f. Hammer et al., 1998; Linnemann and Romer, 2002), whereas types I and IIa are scattered across the field of within-plate tholeiites (WPT) and volcanic arc basalts. In particular, major element values of type I intercalations (Table 10 and Figs. 9 and 10B) point to mixing with sediment (graywacke). But there is also a strong similarity in element composition (and pattern of variation) with the metabasic
0.17 9.53 0.03 51.74 30.23 0.30 2.74 2.53 0.04 0.06 97.36
Na2O K2O CaO SiO2 Al2O3 TiO2 FeO MgO MnO Cr2O3 Sum
0.02 0.79 0.00 0.81 3.35 0.65 1.66 2.31 0.01 0.00 0.15 0.24 0.00
6.87
Na K Ca Sum alk Si AlIV AlVI Al tot. Ti Cr3+ Fe 2+ Mg Mn
Total cations
11 Ox:
III 0 877b
Region Zone Sample
White mica
6.85
0.01 0.71 0.01 0.73 3.31 0.69 1.66 2.35 0.01 0.00 0.22 0.23 0.00
0.09 8.62 0.08 51.07 30.71 0.24 4.13 2.43 0.00 0.00 97.37
III 0 877b
6.97
0.08 0.86 0.00 0.94 3.08 0.92 1.86 2.78 0.03 0.00 0.08 0.06 0.00
0.66 10.37 0.00 47.61 36.50 0.55 1.48 0.61 0.00 0.02 97.79
III 0 877b
6.98
0.02 0.90 0.00 0.92 3.17 0.83 1.60 2.43 0.09 0.00 0.21 0.16 0.00
0.19 10.78 0.00 48.27 31.37 1.81 3.84 1.68 0.02 0.00 97.97
III 0 877b
7.03
0.04 0.89 0.00 0.93 3.10 0.90 1.69 2.59 0.04 0.00 0.26 0.11 0.00
0.31 10.48 0.00 46.54 32.97 0.81 4.60 1.07 0.08 0.00 96.85
III 0 877b
6.98
0.05 0.82 0.00 0.87 3.11 0.89 1.71 2.59 0.04 0.00 0.26 0.11 0.00
0.38 9.54 0.03 46.43 32.82 0.73 4.61 1.15 0.09 0.00 95.77
III 0 877b
6.94
0.12 0.79 0.00 0.91 3.16 0.84 1.78 2.63 0.05 0.00 0.08 0.11 0.00
0.97 9.59 0.00 48.78 34.43 1.01 1.44 1.17 0.00 0.06 97.45
III 0 877b
6.94
0.10 0.83 0.00 0.93 3.13 0.87 1.78 2.65 0.06 0.00 0.07 0.10 0.00
0.79 10.04 0.00 48.31 34.74 1.25 1.37 1.08 0.01 0.05 97.64
III 0 877b
6.99
0.04 0.88 0.00 0.92 3.28 0.72 1.61 2.33 0.02 0.00 0.24 0.20 0.00
0.31 10.47 0.00 49.79 30.02 0.42 4.40 2.05 0.04 0.02 97.52
III 0 877b
6.94
0.01 0.79 0.01 0.81 3.38 0.62 1.50 2.12 0.01 0.00 0.28 0.34 0.00
0.09 9.39 0.09 51.26 27.26 0.17 5.05 3.49 0.04 0.06 96.90
III 0 877b
6.96
0.09 0.76 0.00 0.85 3.09 0.91 1.75 2.66 0.04 0.00 0.22 0.10 0.00
0.69 9.05 0.05 46.85 34.26 0.80 3.92 1.00 0.01 0.02 96.64
III 0 877b
6.93
0.13 0.79 0.00 0.92 3.09 0.91 1.86 2.76 0.04 0.00 0.05 0.07 0.00
1.02 9.56 0.02 47.56 36.04 0.78 1.01 0.73 0.00 0.00 96.72
III 0 877b
7.02
0.04 0.89 0.00 0.93 3.21 0.79 1.59 2.38 0.05 0.00 0.28 0.16 0.01
0.31 10.50 0.00 48.45 30.52 1.02 5.12 1.62 0.17 0.00 97.70
III 0 877b
6.95
0.09 0.83 0.00 0.92 3.10 0.90 1.86 2.76 0.04 0.00 0.07 0.06 0.00
0.72 10.03 0.00 47.98 36.27 0.76 1.24 0.60 0.02 0.01 97.62
III 0 877b
6.98
0.05 0.90 0.00 0.95 3.11 0.89 1.85 2.74 0.00 0.00 0.06 0.12 0.00
0.42 10.76 0.00 47.68 35.55 0.08 1.18 1.20 0.06 0.01 96.93
III 0 877b
6.97
0.03 0.84 0.00 0.87 3.13 0.87 1.78 2.65 0.01 0.00 0.15 0.16 0.00
0.27 10.06 0.00 47.93 34.43 0.25 2.66 1.63 0.00 0.00 97.23
I 2 777a
TABLE 6. CHEMICAL ANALYSES OF DETRITAL WHITE MICA BASED ON MICROPROBE (WDS) DATA
6.93
0.06 0.84 0.00 0.90 3.20 0.80 1.72 2.52 0.06 0.00 0.10 0.15 0.00
0.48 10.02 0.02 48.95 32.79 1.18 1.82 1.56 0.03 0.04 96.88
I 2 907
6.88
0.05 0.79 0.00 0.84 3.21 0.79 1.75 2.54 0.06 0.00 0.09 0.14 0.00
0.39 9.41 0.02 49.02 32.83 1.26 1.62 1.47 0.01 0.00 96.02
I 2 907
6.93
0.05 0.79 0.00 0.84 3.17 0.83 1.73 2.57 0.03 0.00 0.15 0.17 0.00
0.38 9.28 0.00 47.57 32.74 0.69 2.78 1.75 0.03 0.00 95.22
I 2 907
Continued
6.89
0.07 0.76 0.00 0.83 3.11 0.89 1.86 2.75 0.03 0.00 0.09 0.08 0.00
0.52 8.97 0.00 46.56 35.01 0.70 1.64 0.78 0.00 0.05 94.22
I 2 907
114 Kemnitz
0.51 9.87 0.00 47.75 35.01 0.98 2.14 1.03 0.02 0.00 97.29
Na2O K2O CaO SiO2 Al2O3 TiO2 FeO MgO MnO Cr2O3 Sum
0.06 0.82 0.00 0.88 3.11 0.89 1.79 2.69 0.05 0.00 0.12 0.10 0.00
6.95
Na K Ca Sum alk Si AlIV AlVI Al tot. Ti Cr3+ Fe2+ Mg Mn
Total cations
11 Ox:
I 2 907
Region Zone Sample
White mica
6.94
0.06 0.84 0.00 0.90 3.14 0.86 1.75 2.60 0.06 0.00 0.12 0.12 0.00
0.50 10.00 0.00 47.92 33.67 1.23 2.10 1.19 0.03 0.00 96.64
I 2 907
6.94
0.09 0.71 0.00 0.80 3.11 0.89 1.75 2.64 0.02 0.00 0.22 0.15 0.00
0.71 8.61 0.02 47.88 34.41 0.36 3.97 1.59 0.05 0.01 97.61
III 1–2 916
6.85
0.04 0.73 0.00 0.77 3.19 0.81 1.81 2.62 0.05 0.00 0.09 0.13 0.00
0.34 8.83 0.00 49.40 34.38 0.95 1.69 1.32 0.00 0.00 96.90
III 1–2 916
6.83
0.03 0.72 0.00 0.75 3.29 0.71 1.77 2.48 0.02 0.00 0.12 0.17 0.00
0.21 8.83 0.03 51.36 32.76 0.38 2.17 1.80 0.01 0.03 97.58
III 1–2 916
6.90
0.05 0.74 0.00 0.79 3.30 0.70 1.66 2.36 0.02 0.00 0.24 0.19 0.00
0.36 8.86 0.02 50.19 30.42 0.42 4.30 1.96 0.00 0.00 96.53
III 1–2 916
6.79
0.02 0.69 0.00 0.71 3.28 0.72 1.70 2.42 0.06 0.00 0.10 0.22 0.00
0.18 8.54 0.01 51.54 32.26 1.31 1.79 2.36 0.00 0.01 98.01
III 1–2 916
6.89
0.04 0.79 0.00 0.83 3.18 0.82 1.82 2.65 0.03 0.00 0.11 0.09 0.00
0.32 9.67 0.01 49.70 35.14 0.65 2.02 0.92 0.04 0.00 98.45
III 1–2 916
6.85
0.08 0.71 0.00 0.79 3.10 0.90 1.89 2.79 0.04 0.00 0.07 0.06 0.00
0.67 8.67 0.01 48.25 36.80 0.79 1.39 0.67 0.01 0.06 97.33
III 1–2 916
6.91
0.02 0.77 0.00 0.79 3.22 0.78 1.66 2.45 0.06 0.00 0.24 0.15 0.00
0.15 9.20 0.02 48.90 31.52 1.14 4.30 1.53 0.01 0.00 96.77
III 1–2 916
6.89
0.12 0.71 0.00 0.83 3.15 0.85 1.81 2.66 0.03 0.00 0.09 0.13 0.00
0.94 8.36 0.02 47.05 33.77 0.67 1.64 1.35 0.00 0.04 93.84
III 1–2 916
6.90
0.05 0.79 0.00 0.84 3.14 0.86 1.82 2.67 0.03 0.00 0.13 0.09 0.00
0.43 9.56 0.00 48.33 34.85 0.67 2.36 0.89 0.00 0.00 97.09
III 1–2 916
6.89
0.04 0.76 0.00 0.80 3.19 0.81 1.73 2.54 0.04 0.00 0.21 0.11 0.00
0.32 9.12 0.00 48.84 32.96 0.91 3.81 1.17 0.00 0.01 97.13
III 1–2 916
6.92
0.11 0.72 0.00 0.83 3.10 0.90 1.81 2.71 0.04 0.00 0.16 0.08 0.00
0.89 8.54 0.00 46.55 34.54 0.76 2.84 0.78 0.00 0.07 94.96
III 1–2 916
6.86
0.04 0.76 0.00 0.80 3.13 0.87 1.85 2.72 0.05 0.00 0.08 0.08 0.00
0.36 9.34 0.02 49.23 36.22 0.97 1.47 0.80 0.08 0.00 98.48
III 1 945
6.93
0.06 0.82 0.00 0.88 3.12 0.88 1.89 2.76 0.00 0.00 0.15 0.02 0.00
0.45 10.03 0.01 48.93 36.75 0.04 2.74 0.16 0.01 0.04 99.16
III 1 945
TABLE 6. CHEMICAL ANALYSES OF DETRITAL WHITE MICA BASED ON MICROPROBE (WDS) DATA (continued)
6.86
0.08 0.71 0.00 0.79 3.13 0.87 1.86 2.74 0.05 0.00 0.08 0.07 0.00
0.63 8.68 0.00 48.83 36.21 1.01 1.51 0.69 0.00 0.02 97.58
III 1 945
6.91
0.05 0.81 0.00 0.86 3.12 0.88 1.82 2.71 0.06 0.00 0.08 0.08 0.00
0.43 9.92 0.00 48.82 35.97 1.23 1.41 0.80 0.00 0.05 98.64
III 1 945
6.91
0.06 0.81 0.00 0.87 3.11 0.89 1.86 2.75 0.04 0.00 0.06 0.08 0.00
0.48 9.90 0.01 48.80 36.54 0.84 1.04 0.81 0.07 0.02 98.50
III 1 945
6.90
0.07 0.79 0.00 0.86 3.12 0.88 1.94 2.82 0.00 0.00 0.09 0.01 0.00
0.55 9.55 0.03 48.26 37.00 0.00 1.65 0.12 0.00 0.05 97.20
III 1 945
The Lausitz graywackes, Saxo-Thuringia, Germany 115
Na K Ca Sum alk Si AlIV AlVI Al total Ti Cr3+ Fe2+ Mg Mn Total cation
0.01 0.85 0.00 0.86 2.69 1.31 0.36 1.67 0.13 0.00 1.33 1.08 0.01 7.77
0.10 8.64 0.02 35.07 18.43 2.25 20.78 9.41 0.14 0.04 94.88
Na2O K2O CaO SiO2 Al2O3 TiO2 FeO MgO MnO Cr2O3 Sum
11 Ox:
I 2 777a
Region Zone Sample
Biotite
0.00 0.83 0.00 0.84 2.66 1.34 0.40 1.74 0.11 0.00 1.36 1.05 0.01 7.77
0.03 8.42 0.02 34.30 18.97 1.95 20.87 9.11 0.13 0.03 93.84
I 2 777a
0.01 0.88 0.00 0.89 2.73 1.27 0.39 1.66 0.12 0.00 1.29 1.07 0.01 7.77
0.06 9.00 0.00 35.59 18.36 2.03 20.16 9.38 0.13 0.03 94.74
I 2 777a
0.01 0.90 0.00 0.91 2.70 1.30 0.36 1.66 0.16 0.00 1.32 1.00 0.01 7.76
0.04 9.08 0.01 34.79 18.13 2.74 20.34 8.60 0.18 0.00 93.91
I 2 777a
0.01 0.88 0.00 0.89 2.72 1.28 0.39 1.67 0.15 0.00 1.33 0.96 0.01 7.74
0.04 8.92 0.02 35.03 18.17 2.59 20.50 8.28 0.16 0.03 93.73
I 2 777a
0.02 0.89 0.00 0.92 2.63 1.37 0.42 1.80 0.15 0.00 1.30 0.97 0.02 7.78
0.15 9.01 0.02 33.78 19.63 2.57 19.99 8.34 0.33 0.02 93.85
I 2 907
0.02 0.89 0.00 0.91 2.65 1.35 0.43 1.78 0.15 0.00 1.26 0.98 0.02 7.76
0.14 8.93 0.00 33.92 19.31 2.54 19.32 8.44 0.26 0.03 92.89
I 2 907
0.02 0.88 0.00 0.90 2.73 1.27 0.47 1.74 0.14 0.00 1.20 0.98 0.02 7.71
0.12 8.90 0.00 35.04 18.96 2.34 18.34 8.46 0.24 0.03 92.44
I 2 907
0.02 0.88 0.00 0.90 2.68 1.32 0.43 1.75 0.13 0.00 1.26 1.00 0.02 7.75
0.11 8.83 0.00 34.29 19.02 2.30 19.29 8.62 0.31 0.04 92.79
I 2 907
0.01 0.86 0.00 0.87 2.64 1.36 0.42 1.78 0.14 0.00 1.32 1.00 0.02 7.77
0.08 8.59 0.00 33.77 19.30 2.31 20.14 8.56 0.29 0.00 93.04
I 2 907
0.02 0.87 0.00 0.89 2.70 1.30 0.44 1.74 0.15 0.00 1.25 0.98 0.02 7.72
0.11 8.75 0.02 34.59 18.94 2.55 19.08 8.45 0.24 0.00 92.73
I 2 907
0.02 0.88 0.00 0.89 2.70 1.30 0.41 1.71 0.15 0.00 1.29 0.97 0.02 7.74
0.10 8.81 0.01 34.56 18.63 2.62 19.74 8.39 0.26 0.02 93.15
I 2 907
TABLE 7. CHEMICAL ANALYSES OF DETRITAL BIOTITE BASED ON MICROPROBE (WDS) DATA
0.00 0.91 0.00 0.91 2.65 1.35 0.43 1.78 0.14 0.00 1.32 0.96 0.02 7.78
0.02 9.15 0.00 33.95 19.40 2.33 20.22 8.28 0.27 0.00 93.62
I 2 907
0.01 0.90 0.00 0.91 2.68 1.32 0.43 1.76 0.14 0.00 1.27 0.99 0.02 7.76
0.08 9.09 0.00 34.46 19.19 2.38 19.52 8.56 0.29 0.04 93.59
I 2 907
0.01 0.83 0.00 0.83 2.85 1.15 0.39 1.54 0.11 0.01 1.29 1.04 0.02 7.68
0.04 8.56 0.02 37.68 17.27 1.93 20.47 9.20 0.26 0.18 95.60
III 1 945
0.00 0.83 0.00 0.83 2.85 1.15 0.39 1.54 0.10 0.01 1.33 1.01 0.02 7.69
0.02 8.53 0.04 37.45 17.19 1.78 20.87 8.95 0.28 0.10 95.22
III 1 945
116 Kemnitz
L B
E
II
RB
Ib
Ia
I
CZECH REPUBLIC
IV
v
v
v
v
v
GÖ RL ITZ SY N C LINE
VA L L E Y
Ie
KM
Id
N
v
v
v
v
v
+
v
v
v
++
v
v
v
v
v
v
v
v
v v
v
GÖ
IIIa
v
+ + + + +
+
IIIb
POLAND
ZG
S
M Vb
K QF
Qt
C
Vi
Va
P
QM
Detrital Composition
Figure 5. Average detrital compositions of the Lausitz Group in regions I–IV plotted as pie charts. Ie—conglomeratic enclave in granodiorite. Detrital composition: C—chert; K—potassium feldspar; M—metamorphic rocks; P—plagioclase; QF—quartz-feldspar aggregates; QM—monocrystalline quartz; Qt—quartzite/sandstone; S—sedimentary rocks; Va—acid volcanic rocks; Vb—basic volcanic rocks; Vi—intermediate volcanic rocks. KM—Kamenz; RB—Radeberg; GÖ—Görlitz; ZG—Zgorzelec. The map fills are as in Figure 2. (For further explanation see the text.)
E
Dresden
R.
20 km
Elb e
Ic
3-4
2-3
1-2
0 -1
Contact-metamorphic zones:
The Lausitz graywackes, Saxo-Thuringia, Germany 117
118
Kemnitz
QM
I - Kamenz, incl. NW regions II - Radeberg III - Görlitz
n ge oro
pl pr uton ov ic en -ar an c ce
led
k loc l-b e nta nc ne na nti ve co pro
vo pr lcan ov ic en -ar an c ce
yc
I - Conglomeratic enclave in granodiorite
rec
IV - "Gröditzer Skala"
QM
B
co nti pr nen ov tal en -bl an oc ce k
Lausitz Group, sample regions:
A
plu ton
m
Lm
F
QP
ic
ag
m
vo
at
lca
ic
nic
ar
c
Lt
n
io llis
co n
arc ati c gm
ge
K magmatic arc
n
ma
oro
gi ed ar rift al m nt ne
nti
co
P
Ls
Lv mixed magmatic arc + rifted continental margin
Li
Ls
Figure 6. Detrital compositions of the Lausitz graywackes plotted on discrimination diagrams defining provenance areas (after Dickinson and Suczek, 1979) and depositional settings (after Ingersoll and Suczek, 1979). Calculated conservatively with P = 100% F. F = P + K, K—potassium feldspar; P—plagioclase; in (B) F = P + K + QF, QF—quartz-feldspar and feldspar-feldspar aggregates; Li = Lv + F + QM, Lv—volcanic-rock fragments, QM—monocrystallne quartz; Lm = QM + M, M—metamorphic-rock fragments; Ls = S + QP, Ls—sedimentary fragments; Lm—metamorphic fragments; Lt = Lv + Ls + Lm + QP (lithoclasts total); QP—polycrystalline quartz, including C (chert) and QT (quartzite/sandstone).
lod (sample/C1 chondrite)
100 region, sample: I 659 I 954 Id 867 Id 871 II 692 II 667
10
1
Lausitz Group pelitic intervals
enclaves (type IIa) in anatexites and granodiorites (Figs. 9 and 10A,B). It can be concluded, therefore, that both represent more or less diluted volcanogenic mass flows derived from the same volcanic source area, likely at the same time. The intercalations of magmatic origin also differ from the volcanogenic layers and enclaves (types I and IIa) in their larger size and nature of occurrence. Their element distribution (Kemnitz, 1998; Table 10 and Fig. 9) displays a relatively homogeneous geochemical composition. However, major element ratios (Fig. 10B) show a high level of similarity between the fields of type IIa metabasites (volcanogenic enclaves) and those of type IIb. Hence, it is possible that the source rocks are identical to the metabasaltic type IIb. This identity is also suggested by the enclave nature of both types IIa and IIb and their occurrence in anatectic rocks of comparable structural level.
0.1 La Ce Pr
Nd Sm Eu Gd
Tb Dy Ho Er
Tm Yb Lu
Figure 7. Chondrite normalized rare earth element pattern for pelitic intervals within the Lausitz Group.
Detrital Micas Detrital micas partly amount to up to 10% of the mode, with white micas dominating. Their compositions (Fig. 11) differ only slightly, which is especially apparent in Figure 11B (classification after Foster, 1960), in which the white mica analyses
The Lausitz graywackes, Saxo-Thuringia, Germany 0.4
100
A
B
0.3
A
10
K2O/Na2O
Al2O3/SiO2
119
B
0.2
C
0.1
D
A B C D
Ac
tiv
4
6
8
10
12
160
C
c
Oceanic island arc
oceanic island arc continental island arc active continental margin passive continental margin
0.1
SiO2
14
90
80
70
60
50 2
l ar
e
1
0 0
Pa conssive tin ent a
/wt. %/
Fe2O3 + MgO
A A
140 120
V /ppm/
70
D
60
A
Ti/Zr
50
B B
100
C 80
Figure 8. Geochemical classification diagrams (Bhatia, 1983; Bhatia and Crook, 1986; Roser and Korsch, 1986) used to define the basin setting of the Lausitz Group. For comparison, the position of the Katzhütte Group (shaded field) is plotted. Data from Kemnitz and Budzinski (1994), Kemnitz et al. (2002) in (A) and (B).
60 40
40
20
30
DD
0 20
C
B
0
10
5
15
20
25
30
Sc /ppm/
10
D
0 0
2
4
6
8
10
12
La/Sc
of a noncontact metamorphic sample (877b) are split into two groups, muscovites and phengites. This split is interpreted to reflect different source rock provenances: a magmatic source of higher formation temperature for the muscovites and a metamorphic source, possibly formed under pressure-dominated conditions, for the phengites (Guidotti and Sassi, 1976; Massonne and Schreyer, 1987). Analyses from contact-metamorphosed samples instead show a transitional behavior—which means they tend to metamorphic homogeneization—and a muscovite composition. For non- and very-low-grade contact-metamorphic samples, the muscovite position should still represent their primary magmatic composition. The detrital biotites are closely grouped in Figure 11B, which suggests a rather homogeneous source-rock area. However, Figure 11A (after Bea, 1980) shows a distinct separation of analyses from very-low-grade samples and those of more strongly contactmorphosed ones. As with the white micas, this separation implies the preservation of primarily magmatic biotites. The classification diagram of Brigatti and Gregnanin (1987) confirms the derivation of the detrital biotites from granitoid rocks (Fig. 11C).
Aspects of Lithostratigraphic Correlation Subdivision of the Lausitz Group? There are two indications that may define base and top positions: The first relates to the volcanogenic layers (type I), which frequently occur around and south of Kamenz (I) and in regions II and IV (Fig. 2). To the northeast, occurrences and layer thickness decrease (quarry Oßling). In the east (region III, Görlitz–IzeraKarkonosze surroundings), the layers are completely absent. Layer thinning is not only controlled by the distance from the source region, but also by time. The layers are present in anatexites and included in granodiorites (type IIa) that occupy the western and central part of the Lausitz region, caused by post-Cadomian uplift, and represent the oldest Proterozoic rocks. As indicated by their geochemical patterns, types I and IIa are assumed to be of the same provenance, but may represent different structural levels. Hence, graywackes of the Lausitz Group farthest from the anatectic thermal front with few such layers likely represent the younger basin sediments. Geographically, younging direction is to the north, northeast, and east. This direction is in agreement
120
Kemnitz
TABLE 8. CHEMICAL ANALYSES OF DETRITAL FELDSPAR Feldspar Region Zone Sample Mineral SiO2 TiO2 Al2O3 CaO FeO BaO Na2O K2O Sum Si Ti Al Ca Fe Ba Na K Cations total AN AB OR
III 1–2 916 Albite 68.04 0.08 19.76 0.73 0.00 0.07 10.24 0.44 99.36 2.988 0.003 1.023 0.034 0.000 0.001 0.872 0.025 4.946 4 94 3
III 1–2 916 Albite 68.26 0.11 18.99 0.31 0.00 0.00 10.51 0.08 98.26 3.020 0.004 0.990 0.015 0.000 0.000 0.902 0.005 4.936 2 98 1
III 1–2 916 Albite 68.67 0.08 18.75 0.16 0.00 0.07 10.48 0.00 98.21 3.036 0.003 0.977 0.007 0.000 0.001 0.898 0.000 4.922 1 99 0
III 1–2 916 Albite 68.08 0.00 18.64 0.34 0.00 0.00 10.38 0.20 97.64 3.031 0.000 0.978 0.016 0.000 0.000 0.896 0.012 4.933 2 97 1
III 1–2 916 Albite 68.75 0.00 19.52 0.28 0.06 0.00 10.57 0.07 99.25 3.011 0.000 1.008 0.013 0.002 0.000 0.898 0.004 4.936 1 98 0
III 1–2 916 Albite 68.91 0.00 19.04 0.21 0.00 0.00 10.15 0.10 98.41 3.036 0.000 0.988 0.010 0.000 0.000 0.867 0.006 4.907 1 98 1
III 1–2 916 Albite 69.53 0.00 19.29 0.37 0.07 0.00 10.83 0.21 100.3 3.019 0.000 0.987 0.017 0.002 0.000 0.911 0.012 4.948 2 97 1
Feldspar Region Zone Sample Mineral
III 1–2 916 Albite
III 1–2 916 Albite
III 1–2 916 Albite
I 2 777a Albite
I 2 777a Albite
I 2 777a Oligoclase
I 2 777a Albite
SiO2 TiO2 Al2O3 CaO FeO BaO Na2O K2O Sum Si Ti Al Ca Fe Ba Na K Cations total AN
68.71 0.00 19.17 0.26 0.08 0.00 10.10 0.13 98.45 3.028 0.000 0.996 0.012 0.003 0.000 0.863 0.008 4.910 1
68.24 0.00 19.08 0.28 0.13 0.00 10.59 0.17 98.49 3.016 0.000 0.994 0.013 0.005 0.000 0.908 0.010 4.946 1
68.70 0.00 18.89 0.19 0.05 0.00 10.25 0.27 98.35 3.034 0.000 0.983 0.009 0.002 0.000 0.878 0.015 4.921 1
68.30 0.00 19.82 0.85 0.00 0.00 9.68 0.15 98.8 3.002 0.000 1.026 0.040 0.000 0.000 0.825 0.009 4.902 5
68.56 0.00 18.77 0.32 0.00 0.00 10.67 0.05 98.37 3.030 0.000 0.978 0.015 0.000 0.000 0.914 0.003 4.940 2
66.51 0.00 21.02 2.22 0.00 0.00 9.97 0.01 99.73 2.920 0.000 1.087 0.105 0.000 0.000 0.848 0.001 4.961 11
67.92 0.00 19.33 0.75 0.00 0.00 10.16 0.13 98.29 3.006 0.000 1.008 0.036 0.000 0.000 0.871 0.007 4.928 4
AB OR
98 1
98 1
97 2
94 1
98 0
89 0
95 1 Continued
The Lausitz graywackes, Saxo-Thuringia, Germany
121
TABLE 8. CHEMICAL ANALYSES OF DETRITAL FELDSPAR (continued) Feldspar Region Zone Sample Mineral SiO2 TiO2 Al2O3 CaO FeO BaO Na2O K2O Sum Si Ti Al Ca Fe Ba Na K Cations total AN AB OR
I 2 777a Albite 68.52 0.00 19.06 0.42 0.00 0.00 10.17 0.06 98.23 3.027 0.000 0.992 0.020 0.000 0.000 0.871 0.003 4.913 2 97 0
I 2 777a Oligoclase 67.40 0.00 21.26 2.52 0.00 0.00 9.19 0.03 100.4 2.930 0.000 1.089 0.117 0.000 0.000 0.775 0.002 4.913 13 87 0
III 1–2 945 Orthoclase 66.20 0.09 17.75 0.00 0.00 0.08 0.75 16.33 101.2 3.025 0.003 0.956 0.000 0.000 0.001 0.066 0.952 5.003 0 6 94
III 1–2 945 Orthoclase 64.80 0.00 17.21 0.29 0.00 0.30 0.16 16.81 99.57 3.026 0.000 0.947 0.015 0.000 0.006 0.015 1.001 5.010 1 1 97
III 1–2 945 Orthoclase 65.74 0.00 18.30 0.04 0.43 0.73 0.54 15.74 101.52 3.004 0.000 0.986 0.002 0.016 0.013 0.048 0.917 4.986 0 5 95
III 1–2 945 Albite
III 1–2 945 Orthoclase
68.15 0.00 19.56 0.73 0.26 0.17 10.20 0.15 99.22 2.997 0.000 1.014 0.034 0.009 0.003 0.869 0.008 4.934 4 95 1
63.25 0.20 17.44 0.22 1.15 0.15 0.11 15.89 98.41 2.992 0.007 0.972 0.011 0.045 0.003 0.010 0.959 4.999 1 1 98
III 1–2 945 Orthoclase
III 1–2 945 Orthoclase
Feldspar Region Zone Sample Mineral SiO2 TiO2 Al2O3 CaO FeO BaO Na2O K2O Sum Si Ti Al Ca Fe Ba Na K Cations total AN AB OR
III 1–2 945 Orthoclase 65.90 0.00 18.00 0.13 0.32 0.04 0.25 17.36 102 3.006 0.000 0.968 0.006 0.012 0.001 0.022 1.011 5.026 1 2 97
III 1–2 945 Orthoclase 65.12 0.00 17.43 0.00 0.11 0.18 0.36 16.66 99.86 3.026 0.000 0.954 0.000 0.004 0.003 0.033 0.987 5.007 0 3 97
III 1–2 945 Albite 69.63 0.00 18.60 0.00 0.03 0.00 10.72 0.19 99.17 3.049 0.000 0.960 0.000 0.001 0.000 0.910 0.010 4.930 0 99 1
III 1–2 945 Orthoclase 64.99 0.35 17.77 0.09 0.00 0.25 0.62 16.32 100.39 3.003 0.012 0.968 0.005 0.000 0.005 0.056 0.962 5.011 0 5 94
III 1–2 945 Orthoclase 64.40 0.02 17.80 0.01 0.00 0.38 0.16 16.78 99.55 3.007 0.001 0.980 0.000 0.000 0.007 0.015 1.000 5.010 0 1 99
66.69 0.01 18.08 0.00 0.10 0.42 0.40 16.67 102.37 3.020 0.000 0.965 0.000 0.004 0.008 0.035 0.963 4.995 0 4 96
63.83 0.06 17.18 0.07 0.00 0.22 0.35 16.36 98.07 3.021 0.002 0.958 0.003 0.000 0.004 0.032 0.988 5.008 0 3 97 Continued
122
Kemnitz
TABLE 8. CHEMICAL ANALYSES OF DETRITAL FELDSPAR (continued) Feldspar Region Zone Sample Mineral
I 2 907 Orthoclase
SiO2 TiO2 Al2O3 CaO FeO BaO Na2O K2O Sum Si Ti Al Ca Fe Ba Na K Cations total AN AB OR
63.14 0.00 17.66 0.00 0.00 1.32 0.39 15.38 97.89 3.003 0.000 0.990 0.000 0.000 0.025 0.036 0.933 4.987 0 4 96
I 2 907 Orthoclase 65.13 0.25 17.63 0.00 0.00 0.34 0.30 16.51 100.16 3.016 0.009 0.962 0.000 0.000 0.006 0.027 0.975 4.995 0 3 97
I 2 907 Albite 69.40 0.00 18.76 0.00 0.00 0.53 10.76 0.12 99.57 3.039 0.000 0.968 0.000 0.000 0.009 0.913 0.007 4.936 0 99 1
I 2 907 Orthoclase 66.39 0.00 18.28 0.00 0.00 0.50 0.79 15.69 101.65 3.016 0.000 0.979 0.000 0.000 0.009 0.070 0.909 4.983 0 7 93
I 2 907 Albite 69.54 0.00 19.17 0.23 0.00 0.29 10.55 0.18 99.96 3.029 0.000 0.984 0.011 0.000 0.005 0.891 0.010 4.930 1 98 1
I 2 907 Albite 69.54 0.00 19.78 0.42 0.00 0.00 10.35 0.08 100.17 3.014 0.000 1.010 0.019 0.000 0.000 0.870 0.004 4.917 2 97 0
I 2 907 Orthoclase 64.70 0.25 18.06 0.00 0.00 1.03 0.27 15.39 99.7 3.006 0.009 0.989 0.000 0.000 0.019 0.025 0.912 4.960 0 3 97
Feldspar Region Zone Sample Mineral
I 2 907 Orthoclase
I 2 907 Oligoclase
SiO2 TiO2 Al2O3 CaO FeO BaO Na2O K2O Sum Si Ti Al Ca Fe Ba Na K Cations total AN AB OR
63.69 0.41 17.70 0.12 0.00 0.84 0.38 15.67 98.81 2.995 0.015 0.981 0.006 0.000 0.016 0.034 0.940 4.987 1 3 96
65.56 0.01 21.03 2.49 0.00 0.00 9.23 0.19 98.51 2.913 0.000 1.101 0.118 0.000 0.000 0.795 0.011 4.938 13 86 1
I 2 907 Orthoclase 66.66 0.00 18.61 0.07 0.00 0.70 0.37 15.86 102.27 3.012 0.000 0.991 0.003 0.000 0.012 0.032 0.914 4.964 0 3 96
I 2 907 Orthoclase 65.54 0.01 17.68 0.01 0.00 0.26 0.69 15.95 100.14 3.025 0.000 0.962 0.001 0.000 0.005 0.062 0.939 4.994 0 6 94
I 2 907 Orthoclase 65.11 0.00 17.74 0.00 0.00 0.19 0.44 16.43 99.91 3.018 0.000 0.969 0.000 0.000 0.003 0.040 0.972 5.002 0 4 96
I 2 907 Orthoclase 65.81 0.00 18.25 0.01 0.00 1.10 0.28 16.93 102.38 3.002 0.000 0.981 0.000 0.000 0.020 0.025 0.985 5.013 0 2 98
III 0 877b Orthoclase 64.40 0.00 18.14 0.00 0.00 0.46 0.00 16.09 99.09 3.007 0.000 0.998 0.000 0.000 0.008 0.000 0.959 4.972 0 0 100 Continued
The Lausitz graywackes, Saxo-Thuringia, Germany
123
TABLE 8. CHEMICAL ANALYSES OF DETRITAL FELDSPAR (continued) Feldspar Region Zone Sample Mineral SiO2 TiO2 Al2O3 CaO FeO BaO Na2O K2O Sum Si Ti Al Ca Fe Ba Na K Cations total AN AB OR
III 0 877b Orthoclase 65.84 0.00 17.87 0.14 0.00 0.51 0.18 16.82 101.36 3.018 0.000 0.966 0.007 0.000 0.009 0.016 0.984 5.000 1 2 98
III 0 877b Orthoclase 66.27 0.00 17.89 0.11 0.12 0.24 0.28 16.70 101.61 3.023 0.000 0.962 0.005 0.005 0.004 0.025 0.972 4.996 0 2 97
III 0 877b Orthoclase 65.36 0.00 18.07 0.00 0.19 0.00 0.26 17.17 101.05 3.005 0.000 0.979 0.000 0.007 0.000 0.023 1.007 5.021 0 2 98
III 0 877b Orthoclase
III 0 877b Orthoclase
63.77 0.43 17.72 0.00 0.56 0.57 0.37 15.95 99.37 2.988 0.015 0.979 0.000 0.022 0.010 0.033 0.953 5.000 0 3 97
64.57 0.00 18.07 0.00 0.12 1.07 0.78 15.34 99.95 3.001 0.000 0.990 0.000 0.005 0.019 0.070 0.909 4.994 0 7 93
III 0 877b Orthoclase
I 2 778 Oligoclase
III 0 877b Orthoclase 65.10 0.11 18.05 0.02 0.16 1.19 0.48 15.52 100.63 3.006 0.004 0.982 0.001 0.006 0.021 0.043 0.914 4.977 0 4 95
III 0 877b Orthoclase 64.85 0.00 17.96 0.06 0.06 1.27 0.38 16.09 100.67 3.004 0.000 0.981 0.003 0.002 0.023 0.034 0.951 4.998 0 3 96
Feldspar Region Zone Sample Mineral SiO2 TiO2 Al2O3 CaO FeO BaO Na2O K2O Sum Si Ti Al Ca Fe Ba Na K Cations total AN AB OR
III 0 877b Albite 67.69 0.00 19.50 0.96 0.50 0.15 10.28 0.22 99.3 2.984 0.000 1.013 0.045 0.018 0.003 0.879 0.012 4.954 5 94 1
III 0 877b Orthoclase 63.57 0.00 18.17 0.00 0.34 1.34 0.52 15.99 99.93 2.978 0.000 1.003 0.000 0.013 0.025 0.047 0.955 5.021 0 5 95
III 0 877b Orthoclase 64.48 0.18 17.75 0.00 0.31 1.19 0.36 15.82 100.09 3.003 0.006 0.974 0.000 0.012 0.022 0.032 0.940 4.989 0 3 97
64.79 0.27 18.00 0.15 0.19 0.37 0.45 16.19 100.41 2.996 0.009 0.981 0.007 0.007 0.007 0.040 0.955 5.002 1 4 95
65.54 0.00 21.38 3.06 0.06 0.00 9.49 0.17 99.7 2.889 0.000 1.110 0.144 0.002 0.000 0.811 0.009 4.965 15 84 1
I 2 778 Oligoclase 65.30 0.00 22.12 3.94 0.04 0.00 8.84 0.12 100.36 2.861 0.000 1.142 0.185 0.001 0.000 0.751 0.007 4.947 20 80 1
I 2 778 Albite 67.74 0.00 19.85 1.11 0.13 0.00 10.89 0.06 99.78 2.970 0.000 1.026 0.052 0.005 0.000 0.926 0.003 4.982 5 94 0 Continued
124
Kemnitz
TABLE 8. CHEMICAL ANALYSES OF DETRITAL FELDSPAR (continued) Feldspar Region Zone Sample Mineral SiO2 TiO2 Al2O3 CaO FeO BaO Na2O K2O Sum Si Ti Al Ca Fe Ba Na K Cations total AN AB OR
I 2 778 Oligoclase 67.90 0.00 20.60 2.12 0.02 0.00 9.24 0.13 100.01 2.959 0.000 1.058 0.099 0.001 0.000 0.781 0.007 4.905 11 88 1
I 2 778 Oligoclase 68.23 0.00 21.75 2.84 0.04 0.00 9.35 0.13 102.34 2.917 0.000 1.096 0.130 0.001 0.000 0.775 0.007 4.926 14 85 1
I 2 778 Oligoclase 64.04 0.00 21.37 3.51 0.00 0.00 9.27 0.10 98.29 2.867 0.000 1.128 0.168 0.000 0.000 0.805 0.006 4.974 17 82 1
I 2 778 Oligoclase 66.08 0.00 22.20 3.91 0.03 0.00 8.07 0.10 100.39 2.881 0.000 1.141 0.183 0.001 0.000 0.682 0.005 4.893 21 78 1
I 2 778 Oligoclase 65.98 0.00 22.31 3.86 0.04 0.00 7.96 0.14 100.29 2.879 0.000 1.147 0.181 0.001 0.000 0.673 0.008 4.889 21 78 1
Feldspar Region Zone Sample Mineral
I 2 778 Oligoclase
I 2 778 Oligoclase
SiO2
67.80
66.35
69.29
65.42
0.00 21.71 3.08 0.00 0.00 9.40 0.14 102.13 2.909 0.000 1.098 0.142 0.000 0.000 0.782 0.008 4.939 15 84 1
0.00 21.66 2.99 0.00 0.00 9.25 0.14 100.39 2.897 0.000 1.114 0.140 0.000 0.000 0.783 0.008 4.942 15 84 1
0.00 20.34 1.69 0.02 0.00 9.76 0.12 101.22 2.982 0.000 1.031 0.078 0.001 0.000 0.815 0.007 4.914 9 91 1
0.00 22.07 3.98 0.00 0.00 8.77 0.14 100.38 2.864 0.000 1.139 0.186 0.000 0.000 0.745 0.008 4.942 20 79 1
TiO2 Al2O3 CaO FeO BaO Na2O K2O Sum Si Ti Al Ca Fe Ba Na K Cations total AN AB OR
I 2 778 Albite
I 2 778 Oligoclase
Note: Based partly on microprobe (WDS) and EDS (scanning electron microscope) data.
I 2 778 Oligoclase 65.83 0.00 22.05 3.89 0.03 0.00 8.73 0.14 100.67 2.872 0.000 1.134 0.182 0.001 0.000 0.738 0.008 4.935 20 80 1
I 2 778 Oligoclase 66.17 0.00 22.28 3.89 0.00 0.00 8.23 0.13 100.7 2.878 0.000 1.142 0.181 0.000 0.000 0.694 0.007 4.902 21 79 1
The Lausitz graywackes, Saxo-Thuringia, Germany 100
SiO2 Al2O3 FeOtot MgO CaO Na2O K2 O
wt. %
10
1
Greywackes
increasing contact metamorphism
Granodiorites/ Anatexites
Metabasites
0.1
1.1
1.2 2.1 2.2 3.1 3.2
4
5
6
I
IIa
IIb
Rock groups 100
Greywackes
increasing contact metamorphism
Granodiorites/ Anatexites
Metabasites
wt. %
10
Ba Zn Y Yb Sc Sr Rb Ga
1
0.1
1.1
1.2 2.1 2.2 3.1 3.2
4
5
6
I
IIa
IIb
Rock groups 100
Greywackes
increasing contact metamorphism
Granodiorites/ Anatexites
Metabasites
wt. %
10
Ti Mn Zr Cr V Nb La Ni Co
1
0.1
1.1
1.2 2.1 2.2 3.1 3.2
4
5
6
I
IIa
IIb
Rock groups
Figure 9. Average major and minor element distributions for all Lausitz rock groups (see Table 8).
with general grain size decrease from south to north and northeast (Schwab, 1962). The second indication relates to a conglomeratic enclave in granodiorite (Lobst, 1996), which represents strata from greater structural depth and mirrors the magmatic-volcanic arc source area of the Lausitz graywackes (Fig. 6 and Table 1, Ie; sample locations listed in Table 11). Although a correct classification of polycrystalline quartz pebbles is problematic because of internal recrystallization, it is unlikely that cherts were part of the
125
spectrum of source rocks. This conglomerate may represent an older littoral on-lap facies of the preceding stage of basin opening, which is not otherwise exposed within the Lausitz Group. Cadomian Saxo-Thuringia/Teplá-Barrandian Zone In contrast to the Lausitz region, the early evolution of the Cadomian is well preserved in the Teplá-Barrandian. Tholeiitic to calc-alkaline volcanic rocks are intercalated in an older basin succession, thus allowing a stratigraphic subdivision of the Teplá-Barrandian Neoproterozoic (Dolejš and Kraft, 1998). Here, the Šteˇchovice Group, which forms the upper unit, is a turbiditic graywacke succession with rare intercalated volcanogenic layers similar to the Lausitz Group. However, it overlies conglomeratic channel deposits (Dobrˇíš conglomerate) with volcanic/magmatic-arc–derived pebbles. Bernadová and Cháb (1968) suggested deposition in an active-margin setting. Frequent features of synsedimentary tectonic activity indicate that the graywackes were accumulated during a period of increased basin subsidence and first collision-related movements (Chlupácˇ, 1993). A rhyolithic boulder from the Dobrˇíš horizon and a granite boulder from a Lower Cambrian horizon yielded ages (570 Ma and 594 Ma) closely comparable to the Saxo-Thuringian ages of Cadomian magmatism (Dörr et al., 1992). Hence, deposition of the Šteˇchovice Group occurred at about the same time as that of the Lausitz Group. Apart from different lithologies and facies, Klápová and Hyršl (2000) found a rather distinct provenance and basin setting comparing the composition of conglomerate horizons from the Barrandian with those of the Cadomian Saxo-Thuringian. Intermediate and acid volcanic rocks are generally less common than they are in the Saxo-Thuringian units, and granitoids are rare. The pebble spectrum of the conglomerate is dominated by unstable sedimentary rock fragments, of which graywackes make up more than 75% of the mode (c.f. Fiala, 1948; Röhlich, 1964). In contrast, own lithoclast analyses from graywacke material, which forms the matrix of the Dobrˇ íš conglomeratic layer, and two other locations of the Šteˇchovice Group (Table 1, Fig. 12) show a broad source-rock variety, and hence, a balanced ratio of unstable to stable components. Magmatic-rock derived lithoclasts dominate, followed by graywacke and chert fragments. As illustrated in the example of gravel and pebble-sized samples from the Lausitz Group (Fig. 5, regions Ic and Ie), the ratio of unstable to stable components will increase with increasing grain size. Therefore, with regard to reconstruction of source rock areas and tectonic basin settings, care has to be taken not to overestimate the part of unstable components given in conglomerates. However, there are differences in lithoclast composition that indicate different basin settings of the two Cadomian realms (Fig. 12). DISCUSSION In a regional context, no systematic differences occur between the study regions of the Lausitz area, either across the whole domain or in time, that can be related to source-area changes.
126 Group Sample Region
Kemnitz
TABLE 9. MAJOR AND TRACE ELEMENT ANALYSES OF CADOMIAN SYN- TO POSTCOLLISIONAL ANATEXITES AND GRANODIORITE INTRUSIONS 4 713 I
722 I
727 I
781 II
791 II
795 b II
797 II
804 II
823 II
833 SO
851 SO
852 SO
68.9 0.65 15.6 1.88 2.8 2.1 1.53 2.81 2.7 0.16 0.3 1.6
68.3 0.69 15.7 1.29 3.5 2 1.35 3.41 2.6 0.15 0.4 1.5
72.4 0.6 14 1.22 2.7 1.3 1.4 2.86 3.4 0.12 0.2 0.7
68.6 0.7 15.9 1.76 3.2 1.8 1.3 3.62 2.9 0.16 0.3 0.9
68 0.7 16.1 1.6 3.6 2.2 1.59 3.16 2.7 0.15 0.4 1.2
70.4 0.67 15.3 1.14 3.2 1.8 1.34 2.96 2.5 0.15 0.4 1.6
68.8 0.71 15.8 1.76 3.4 2 1.54 3 2.6 0.16 0.3 1.2
67.3 0.81 15.6 1.45 4.3 2.1 1.26 3.57 2.4 0.16 0.5 1.6
69.4 0.66 15.8 1.6 3.2 1.9 1.52 3.24 3 0.17 0.6 1.3
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
70 0.55 14.5 5.40* 3.7 1.1 1.11 3.16 2 0.13 0.3 1.9
72.6 0.51 13.5 4.50* 3.2 1.4 1.27 2.19 3.3 0.13 0.2 1.7
(wt%) SiO2 TiO2 Al2O3 Fe2O3 FeO MgO CaO K2O Na2O P2O5 H2O+ H2O– (ppm) Zn Ba Nb Zr Y Sr Rb Mn Ti B Be Co Cr Cu Ga La Ni Pb Sc Sn V Yb Li
Group Sample Region
63 644 14 202 34 213 96 406 3720 5.6 2 8.9 54.6 75.4 12.7 39.3 26 24 7.7 2.7 76.1 3.4 23
65 1280 14 221 27 178 133 469 4120 12 2 7 86 8 15 50 16 18 12 4 82 2 46
66 723 14 201 21 246 104 221 3670 2.2 0.7 10.4 57 6.5 10.7 36.3 25.2 18.5 10.9 2.2 56.5 2.5 23
65 904 14 231 29 214 111 393 3860 16.5 1.8 10 72.5 7.8 14.5 34.5 22.4 17.6 9.4 3.1 80 3.1 n.d.
83 796 13 205 27 236 121 490 4290 5.3 1.7 13.3 70.7 8 18.6 45.3 43 18.1 12.9 3.5 96.2 2.9 45
77 777 14 230 28 219 108 374 3850 17.1 2.5 16.3 75.1 14.8 13.9 41.7 50.1 12.2 13.9 3.5 85.2 2.9 39
65 696 14 225 30 202 108 414 4330 6.8 1.6 11.7 110.3 19.5 16.3 47 33 16.7 13 3 114.8 4.5 42
85 735 16 240 29 173 141 525 5100 48 3 8 115 34 17 58 24 13 16 4 108 2 46
74 780 14 217 29 196 132 416 4080 7 2 9.8 61.5 17.5 9.3 65.3 24.2 14.3 12.6 2.8 73.6 2.6 46
81 780 13 201 33 181 104 450 4140 20 2.5 9 83 26 n.d. n.d. 25 11 n.d. 9 66 n.d. n.d.
81 976 13 186 21 154 136 590 4320 18 2.8 9 81 10 n.d. n.d. 20 10 n.d. 7 70 n.d. n.d.
84 550 10 229 25 161 96 385 3730 44 3.1 9 66 31 n.d. n.d. 20 18 n.d. 8 61 n.d. n.d.
5 721 I
723 I
726 I
734 I
737 I
724 I
754 I
775 I
802 II
740 BW
817 BA
839 SO
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
68.4 0.73 16 1.55 3.6 1.9 1.51 3.02 2.6 0.12 0.5 1.5
69.2 0.64 15.3 1.71 3.1 2 0.82 2.76 2.3 0.13 n.d. n.d.
71.7 0.51 13.7 4.00* 2.8 1.2 1.79 2.29 3 0.18 0.4 1.5
(wt%) SiO2 TiO2 Al2O3 Fe2O3 FeO MgO CaO K2O Na2O P2O5 H2O+ H2O–
Continued
The Lausitz graywackes, Saxo-Thuringia, Germany
TABLE 9. MAJOR AND TRACE ELEMENT ANALYSES OF CADOMIAN SYN- TO POSTCOLLISIONAL ANATEXITES AND GRANODIORITE INTRUSIONS (continued) Group Sample Region
5 721 I
723 I
726 I
734 I
737 I
724 I
754 I
775 I
802 II
(ppm) Zn Ba Nb Zr Y Sr Rb Mn Ti B Be Co Cr Cu Ga La Ni Pb Sc Sn V Yb Li
57 875 15 228 25 217 100 389 3970 5.2 2.1 6.5 74.7 12.4 18.3 47.5 14.5 20.2 10.2 3.2 85.2 2.6 39
75 749 14 190 26 200 113 441 3810 2.2 1.8 10.4 72.3 16.2 14.2 60.4 27.4 16.7 9.7 2.5 78 2.9 40
55 862 15 212 25 230 102 327 3000 4.2 0.7 7.3 55.5 7.6 16.2 30.5 17.9 24 7.3 3.4 51.9 2.3 22
57 652 16 353 31 256 111 316 3820 10.2 0.7 7.8 65.7 11.9 16 53 20.6 22.4 8.7 4 59.5 2.8 30
75 527 12 236 38 199 96 497 4560 3.3 1.4 13.3 47.6 12.4 15 38.8 25.7 16.9 13 3.1 69.9 3.9 35
61 700 14 241 33 254 96 440 4140 1.9 2.1 10.9 91.2 13.4 15.8 33.2 33.3 16.3 10.4 3 76.5 2.6 31
80 808 12 239 39 217 90 699 5370 2.6 1.7 11.3 58 10.7 14 20.2 20.9 13.1 18.5 2.8 82.2 3.4 35
80 660 12 220 39 203 88 780 8130 1.4 1.9 13 23.2 9.7 20.1 25.4 7.4 12.1 26 4.9 113.5 3.7 27
60 715 14 202 27 188 117 354 3640 6.9 1.9 7 58.1 47.9 12.8 43.5 20.2 15.4 7.7 3.4 60.9 2.1 29
Group Sample Region
6 719 I
725 I
729 I
732 I
735 I
736 I
738 I
752 I
756 I
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
62 711 17 208 35 171 134 264 3590 3.2 0.7 6 13.7 4.8 17.5
70 776 15 274 42 177 120 485 5070 1.6 1.6 9.7 41.6 14.2 17.5
78 724 23 323 44 243 122 318 6090 34 0.5 8 39 9 24
80 790 18 316 34 190 118 547 5100 6.7 2 10 27.9 14.5 21.9
75 728 18 274 38 169 138 552 5330 2.3 3 9.6 38.4 12.3 20.5
60 785 15 262 37 185 97 513 4640 4.7 2.3 9.8 35.5 11.4 18.9
34 805 12 186 35 187 108 338 2740 2.6 2.8 5.6 15.4 4.6 19
64 795 15 233 38 164 137 466 4210 0.4 2.1 9 24 9.7 15.3
740 BW
127
817 BA
839 SO
58 749 18 212 35 102 98 350 3930 9.7 2 5.6 55.8 13.5 10.9 79 30.9 7.1 12.7 3.4 75.5 3 28
63 631 11 236 25 220 108 381 3870 21 2.4 6 67 5 n.d. n.d. 16 16 n.d. 9 51 n.d. n.d.
809 I
813 I
814 I
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
68 793 16 244 38 156 131 468 4350 3 2.5 9.1 25.7 7.9 20.6
71 697 16 259 36 189 115 505 4890 2.7 2.9 10.3 21 11.7 16.8
67 745 16 262 37 161 125 453 4230 0.8 1.5 8.9 21.7 11.4 12.8
55 595 11 182 32 148 127 388 2840 6.6 3.6 8.1 25.6 42.4 11.6
75 780 15 216 31 201 120 455 4320 2 2 10 92 20 18 48 31 20 14 4 85 2 45
(wt%) SiO2 TiO2 Al2O3 Fe2O3 FeO MgO CaO K2O Na2O P2O5 H2O+ H2O– (ppm) Zn Ba Nb Zr Y Sr Rb Mn Ti B Be Co Cr Cu Ga
Continued
128
Kemnitz
TABLE 9. MAJOR AND TRACE ELEMENT ANALYSES OF CADOMIAN SYN- TO POSTCOLLISIONAL ANATEXITES AND GRANODIORITE INTRUSIONS (continued)
Group Sample Region
6 719 I
725 I
28.7 11 16 7.9 3.4 35.9 2.8 28
43.7 14.7 16.4 13.8 3.3 75.3 4.3 42
6 822 II
739 BW
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
65.1 0.96 17 1.18 3.8 2 3.62 3.01 3.5 0.22 0.3 0.7
729 I
732 I
735 I
736 I
738 I
752 I
756 I
809 I
813 I
814 I
56.7 13.1 15.6 14.4 4.1 52.4 2.2 37
46.4 16 18.5 17 4.7 67.4 3.5 52
37.2 15.9 15.8 13.1 3.8 67.3 3.2 42
28.3 12.2 20.9 7.8 3.7 35 3.4 20
47.4 14.9 15.6 12.3 2.9 53.9 3.7 55
38.6 11.9 18.4 12.8 3.5 54.4 3.6 47
60.5 18 16.3 12.3 4.5 50.7 2.7 46
44 13.3 14.7 11.7 3.6 44.5 3.2 48
35.8 21 17.4 8.5 3.5 39.6 2.8 47
819 CZ
829 HW
836 SO
844 SO
65.7 0.84 17.1 1.85 3.1 1.9 2.99 3.09 3.5 0.19 0.4 1.1
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
69.9 0.55 14.7 4.30* 3 1 2.41 2.63 3.6 0.24 0.3 1.4
148 54 18 197 56 308 55 1560 2670 23 7 7 47 73 n.d. n.d. 27 10 n.d. 7 43 n.d. n.d.
67 678 12 218 31 222 93 313 4090 14 2.1 4.5 40 20 n.d. n.d. 9 20 n.d. 9 64 n.d. n.d.
(ppm) La Ni Pb Sc Sn V Yb Li
Group Sample Region
64 16 20 100 2 72 3 45
(wt%) SiO2 TiO2 Al2O3 Fe2O3 FeO MgO CaO K2O Na2O P2O5 H2O+ H2O– (ppm) Zn Ba Nb Zr Y Sr Rb Mn Ti B Be Co Cr Cu Ga La Ni Pb Sc Sn V Yb Li
30 692 14 282 43 206 95 419 4440 6.1 3.4 8.7 34.5 10.2 15.5 42.3 17.2 12.9 12.4 3.8 46.4 3.2 36
83 876 18 333 34 201 110 522 5480 1.1 2.2 12.3 31.1 11.4 21.6 53.6 17.4 15.1 15.2 3.6 63 2.6 33
83 578 18 287 28 221 147 386 4940 7.7 0.5 8.8 36 9.6 13.6 97.1 22.3 15.1 13.2 3.9 70.9 2 30
60 668 12 254 26 159 108 375 3840 26 2.7 7 28 17 n.d. n.d. 9 11 n.d. 8 39 n.d. n.d.
Note: Major elements are in wt.% and trace elements are in ppm. BA—Bautzen; BW—Bischofswerda; CZ—Czorneboh; HW—Hohwald region; SO—Sora. Source: Data from Kemnitz and Budzinski (1994). *FeO total.
The Lausitz graywackes, Saxo-Thuringia, Germany 6
12
A Trachydacite
Trachyandesite
6
K2O/Na2O
8
Trachybasalt Basaltic andesitic Andesite
Basalt
4
B
5
10
K2O + Na2O
129
Tuffaceous greywacke layers Tuffaceous enclaves Metabasaltic enclaves
4
siliciclastica 3
(greywackes)
I
2 Dacite
1
volcanic rocks
2
pyroclastic rocks
IIb
IIa
0
40
45
50
55
SiO2
60
65
70
3
2
1
0
0
4
6
5
7
SiO2/Al2O3
75
/wt. %/ B - E-type MORB
C
C - Within-plate tholeiites and volcanic arc basalts Figure 10. Classification diagrams for metabasic layers within the Lausitz Group and from enclaves in anatectic and granodioritic rocks of the Lausitz area. (A) TAS classification of volcanic rocks after LeBas et al. (1986). (B) K2O/Na2O vs. SiO2/Al2O3 diagram of Wimmenauer (1984). Data from Löffler (1982), Kemnitz and Budzinski (1994), and Nasdala and Pfeiffer (1996). (C) Y-Nb-Zr diagram of Meschede (1986). E-type MORB—enriched mid-oceanic ridge basalt; N-type MORB—normal MORB.
D - N-type MORB and volcanic arc basalts IIa Nb x 2 I
C
B C
Zr/4
The only indication of any changes in detrital input is the higher overall number of detrital micas in study region III (Table 1) compared to region I. In all other cases, a largely uniform source area is indicated. The provenance of the detrital micas and the invariance of the heavy mineral spectrum also indicate a stable and essentially homogeneous source area. This homogeneity corresponds with other data. Apart from an attempt at biostratigraphic subdivision by Burmann (1997, 2001) and the two lithological and sedimentological indications introduced above, there are few lithologic criteria for a lithostratigraphic subdivision. The character of the exposed graywacke series is that of basin sediments that accumulated during continuous long-term subsidence. The detrital component composition and geochemical data show that deposition took place in a subduction-related basin without any compositional change in source-rock supply through time. Neither key beds nor unconformities that might indicate interruptions in this continuous process have been proven to exist. Hence, lithostratigraphic subdivision is not possible. Geochemical classification clearly indicates a subductionrelated basin position in a setting between active-margin and continental island arc, thus supporting the discrimination diagrams based on detrital composition. The main controlling factor affecting the geochemical composition of the studied graywacke members is their source-rock mineral composition.
B
D
D IIb
Y
The external position of the Cadomian Saxo-Thuringian basins within the fragments of peri-Gondwana has been concluded from zircon dating constraining provenance from several source areas, including Armorica (Linnemann et al., 2000). Similar conclusions can be drawn from geochemical and petrographical provenance studies discussed in this article. This significant continental block influence seems to be best explained by a backarc basin setting, as proposed by Buschmann et al. (1995) and Linnemann et al. (2000). However, the marine turbiditic graywacke successions were partly involved in extensive anatectic processes. Assuming forearc basin sediments on-lap on marginal fragments of a disintegrating old continent, with a continentward developed magmatic arc, a comparable scenario would be expected. Transforming this conception into recent geographical positions, subduction would have been directed southward. Crustal heating, initiated above the subducting slab, generated large volumes of anatectic melts beneath the Lausitz graywackes and fully involved the continental island arc. In this context, it would be of interest to determine whether the number of old continental crust–derived zircons exceeds that of volcanic-arc–derived zircons. Folding and anchimetamorphism of the Lausitz Group reflect compressional processes during the Cadomian orogeny. Although the closure of an ocean cannot be assumed (Nance and
96 442 14 236 35 338 112 1290 5440 12.0 2.0 18.0 58.0 97.0 16.0 34.0 38.0 10.0 16.0 3.0 70.0 3.0 26
96 442 14 236 35 338 112 1290 5440 12.0 2.0 18.0 58.0 97.0 16.0 34.0 38.0 10.0 16.0 3.0 70.0 3.0 26
0.026
n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a.
n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a.
(wt%)
SiO2 TiO2 Al2O3 Fe2O3 FeO MnO MgO CaO K2O Na2O P2O5 H2O+ H2O– (ppm) Zn Ba Nb Zr Y Sr Rb Mn Ti B Be Co Cr Cu Ga La Ni Pb Sc Sn V Yb Li Ratios SiO2/Al2O3 K2O/Na2O Zr/TiO2 Na2O + K2O
I 712 I
0.044
9 55 13 197 32 461 9 1720 2710 6.2 2.3 1.7 37.7 13.9 21.2 20.0 5.0 10.7 7.8 3.9 46.8 2.4 5
n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a.
I 745 I
0.065
32 776 11 235 22 161 63 126 2180 13.1 0.7 5.7 26.4 7.4 9.2 30.3 14.0 11.7 5.5 1.9 35.1 1.8 17
n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a.
I 771 I
5.19 0.94 0.040 1.94
84 219 11 190 37 350 36 1050 2820 3.0 2.0 12.0 40.0 18.0 14.0 58.0 28.0 13.0 9.0 4.0 50.0 3.0 15
66.40 0.53 12.80 4.41* 3.97 n.a. 0.90 10.58 0.94 1.00 0.84 0.30 1.60
I 810 I
0.051
46 28 9 224 34 263 8 1820 2640 16.0 4.0 8.0 47.0 36.0 1.0 n.a. 16.0 12.0 n.a. 9.0 48.0 n.a. n.a.
n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a.
I 848 I
0.75
4.92
52 184 11 235 48 279 25 5490 2830 20.0 2.3 9.0 69.0 1.6 n.a. n.a. 21.0 4.9 n.a. 4.8 64.0 n.a. 7
58.00 0.54 11.80 4.15* 3.73 n.a. 1.70 19.77 0.75 0.00 0.33 n.a. n.a.
I 956 I
5.00 0.27 0.170 1.40
n.a. <100 11 170 n.a. n.a. n.a. n.a. n.a. 9.0 2.3 10.0 54.0 26.0 12.0 n.a. 27.0 18.0 n.a. n.a. 66.0 n.a. n.a.
56.00 0.10 11.20 0.10 2.60 0.40 1.30 19.60 0.30 1.10 0.20 n.a. n.a.
I Na2 I
5.38 0.32 0.063 1.24
n.a. 100 12 250 n.a. n.a. n.a. n.a. n.a. 12.0 2.3 10.0 50.0 38.0 14.0 n.a. 28.0 18.0 n.a. n.a. 62.0 n.a. n.a.
64.50 0.40 12.00 0.20 3.20 0.30 1.50 11.30 0.30 0.94 0.20 n.a. n.a.
I Na3 I
0.043
173 284 15 187 35 272 121 1110 2580 31.0 2.0 8.0 32.0 110.0 12.0 45.0 18.0 8.0 8.0 2.0 30.0 3.0 23
n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a.
I 643 II
0.043
67 306 14 170 25 222 50 216 2370 8.8 2.8 12.7 32.5 33.3 16.1 47.3 41.5 16.3 7.3 3.2 34.4 2.7 10
n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a.
I 668 II
0.043
156 67 15 218 36 284 47 1720 3020 8.0 2.3 10.8 75.2 23.7 16.8 63.9 30.7 14.4 10.8 4.7 67.0 3.5 13
n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a.
I 694 II
TABLE 10. MAJOR AND TRACE ELEMENT ANALYSES OF CADOMIAN METABASIC ROCKS OF THE LAUSITZ REGION
I 712 I
Type Sample Region
0.057
22 114 18 238 35 481 12 1820 2940 5.9 2.6 2.8 107.8 27.0 16.2 61.3 6.9 10.1 10.2 4.6 62.0 3.3 4
n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a.
I 704 II
Continued
0.94
4.04 3.70
26 119 14 185 35 569 30 688 3070 7.1 1.7 2.0 46.6 3.6 25.4 36.7 7.5 12.0 10.4 9.5 63.7 2.9 1
67.4 0.56 16.7 4.95* 4.13 n.a. 0.6 12.24 0.74 0.2 0.15 0.3 0.9
I 788 II
130 Kemnitz
Zn Ba Nb Zr Y Sr Rb Mn Ti B Be Co Cr Cu Ga La Ni Pb Sc Sn V Yb Li Ratios SiO2/Al2O3 K2O/Na2O Zr/TiO2 Na2O + K2O
(ppm)
SiO2 TiO2 Al2O3 Fe2O3 FeO MnO MgO CaO K2O Na2O P2O5 H2O+ H2O–
(wt%)
Type Sample Region
4.04 3.7 0.036
26 119 14 185 35 569 30 688 3070 7.1 1.7 2.0 46.6 3.6 25.4 36.7 7.5 12.0 10.4 9.5 63.7 2.9 1
67.4 0.56 16.7 4.59* 4.13 n.a. 0.6 12.24 0.74 0.2 0.15 0.3 0.9
I 788 II
4.56 4.40 0.046 1.08
33 162 15 236 37 295 48 1180 30400 38.0 4.0 4.0 37.0 28.0 14.0 55.0 10.0 10.0 8.0 4.0 42.0 2.0 15
69.3 0.57 15.2 3.76* 3.72 n.a. 1 9.88 0.88 0.2 0.27 0.3 1.1
I 799 II
0.037
71 117 12 190 37 350 42 2280 3120 17.0 2.5 10.5 57.3 17.9 15.6 60.9 40.6 5.0 11.0 3.9 51.4 2.8 10
n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a.
I 765 II
79 248 14 166 36 267 47 1690 3580 13.0 6.0 18.0 70.0 167.0 n.a. n.a. 60.0 12.0 n.a. 7.0 71.0 n.a. n.a.
n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a.
IIa 831 I GD
4.29 1.40 0.040 0.24
42 31 16 187 38 687 11 1760 2800 3.9 1.9 10.9 55.6 13.8 12.6 29.0 28.9 9.5 9.2 3.6 39.9 2.9 4
65.7 0.52 15.3 5.51* 4.6 n.a. 1.2 12.78 0.14 0.1 0.18 0.2 0.5
IIa 805a II AN
4.18 0.58 0.038 0.63
39 56 14 199 39 262 14 1970 3120 41.6 1.9 3.5 44.5 4.3 12.4 65.5 13.9 5.5 11.0 3.0 55.5 3.1 7
66.0 0.61 15.8 4.75* 4.28 n.a. 1.7 13.28 0.23 0.4 0.18 0.3 0.5
IIa 818b AN CZ
76 550 14 274 36 177 109 425 4790 14.0 3.9 10.0 43.0 37.0 n.a. n.a. 16.0 8.0 n.a. 7.0 60.0 n.a. n.a.
n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a.
IIa 828 AN ST
62 446 11 179 28 291 71 979 2880 14.0 4.2 6.0 69.0 31.0 n.a. n.a. 21.0 37.0 n.a. 7.0 43.0 n.a. n.a.
n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a.
IIa 834 AN SO
59 78 12 187 37 423 27 1470 3160 13.0 4.0 6.0 93.0 30.0 n.a. n.a. 23.0 7.0 n.a. 5.0 53.0 n.a. n.a.
n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a.
IIa 837 AN SO
37 111 10 190 29 300 20 1440 2790 15.0 3.2 3.8 57.0 26.0 n.a. n.a. 8.0 9.0 n.a. 8.0 48.0 n.a. n.a.
n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a.
IIa 845 AN SO
27 57 16 203 32 368 29 1710 2980 12.0 2.9 4.5 63.0 33.0 n.a. n.a. 8.0 8.0 n.a. 10.0 44.0 n.a. n.a.
n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a.
IIa 853 AN SO
TABLE 10. MAJOR AND TRACE ELEMENT ANALYSES OF CADOMIAN METABASIC ROCKS OF THE LAUSITZ REGION (continued)
Continued
The Lausitz graywackes, Saxo-Thuringia, Germany 131
46.40 3.12 11.60 n.a. 7.70 n.a. 5.50 9.80 0.24 1.50 0.29 0.80 1.30
(wt%) SiO2 TiO2 Al2O3 Fe2O3 FeO MnO MgO CaO K2O Na2O P2O5 H2O+ H2O– 46.40 3.12 11.60 n.a. 7.70 n.a. 5.50 9.80 0.24 1.50 0.29 0.80 1.30
IIb 808 ANLP
n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. 3.41 0.24 3.23
3.19 0.25 3.00
48.10 2.90 14.10 6.26 7.58 0.24 5.80 10.90 0.63 2.60 0.49 n.a. n.a.
IIb Lö1.2 ANST
n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a.
48.10 3.00 15.10 6.49 6.96 0.27 5.80 10.90 0.60 2.40 0.38 n.a. n.a.
IIb Lö1.1 ANST
2.12
3.64 0.51
n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a.
50.90 2.70 14.00 5.79 7.79 0.23 5.80 10.30 0.72 1.40 0.37 n.a. n.a.
IIb Lö2.1 ANST
2.11
3.65 0.51
n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a.
51.40 2.40 14.10 5.16 8.34 0.23 5.90 9.90 0.71 1.40 0.46 n.a. n.a.
IIb L2.2 ANST
2.4
3.29 1.18
n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a.
46.70 2.50 14.20 5.21 9.17 0.20 7.00 11.80 1.30 1.10 0.82 n.a. n.a.
IIb Lö3.1 ANST
2.4
3.34 1.18
n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a.
46.70 2.50 14.00 5.24 9.24 0.21 6.90 12.00 1.30 1.10 0.81 n.a. n.a.
IIb Lö3.2 ANST
1.72
3.57 0.74
n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a.
47.50 2.30 13.30 5.02 9.52 0.20 6.70 12.90 0.73 0.99 0.84 n.a. n.a.
IIb Lö4.1 ANST
1.74
3.55 0.76
n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a.
47.20 2.40 13.30 5.18 9.44 0.21 6.80 12.90 0.75 0.99 0.83 n.a. n.a.
IIb Lö4.2 ANST
Note: Major elements are in wt% and trace elements in ppm. Type I—tuffaceous graywacke (volcanogenic mass flow members of the Lausitz Group); IIa—volcanogenic mass flow (enclaves in granodiorite and anatexite); IIb—metabasic rock (enclaves in granodiorite [GD] and anatexite [AN]); III—metabasic to intermediate rocks (enclaves in granodiorite and anatexite). CZ—Czorneboh; LP—Leppersdorf region; SO—Sora; ST—Stolpen surroundings. Sources: Data from Löffler (1982) (Lö), Nasdala and Pfeiffer (1991) (Na), and Kemnitz and Budzinski (1994).
Zn n.a. n.a. Ba 271 271 Nb 5 5 Zr 251 251 Y 86 86 Sr 120 120 Rb 9 9 Mn 2480 2480 Ti 19500 19500 B 6.0 6.0 Be 0.3 0.3 Co 28.0 28.0 Cr 51.0 51.0 Cu 58.0 58.0 Ga 16.0 16.0 La 26.0 26.0 Ni 67.0 67.0 Pb 6.0 6.0 Sc 40.0 40.0 Sn 8.0 8.0 V 685.0 685.0 Yb 8.0 8.0 Li 12 12 Ratios SiO2/Al2O3 4.00 4.00 K2O/Na2O 0.16 0.16 Zr/TiO2 0.001 Na2O + K2O 1.74
(ppm)
IIb 808 ANLP
Type Sample Region
TABLE 10. MAJOR AND TRACE ELEMENT ANALYSES OF CADOMIAN METABASIC ROCKS OF THE LAUSITZ REGION (continued)
132 Kemnitz
The Lausitz graywackes, Saxo-Thuringia, Germany 1.00
A
Legend
Detrital biotites
Metamorphic field
A–C:
0.70
0.60
0.50 1.90
2.00
gr ad e
Zone Sample Signature
low
ct- sm ta rphi n co mo a et m
ve ry-
Al VI
0.90
0.80
2.10
133
Region I Id III 2 2 0 1 1-2 777a 907 877b 945 916
Magmatic field
2.30
2.20
2.50
2.40
2.60
Si / Al IV
Legend C: Fields of biotites from
Mg
B ite op Ph
log
Granites Granodiorites Tonalites Diorites
ite Mu
sc
ov
Phe
AlVI+ Ti
C
Nephelinites All other volcanic rocks
h ric g- tite M io b
e ngit
Ti
Detrital biotites of the Lusatian Group
rich Fe- tite o i b biotite
hyllite/ Siderop elane lepidom
Fetot + Mn
evolution trend
AlVI
Mg
Figure 11. Compositions of detrital white micas and biotites in graywackes of the Lausitz Group showing different degrees of contact-metamorphic overprint. (A) AlVI vs. Si/AlIV plot of Bea (1980). (B) Mg-Fe-Al plot modified from Foster (1960). (C) Provenance diagram of Brigatti and Gregnanin (1987).
Murphy, 1994), the geochemical signatures of intercalated basic rocks (type II in Fig. 11) indicate that considerable quantities of oceanic crust were subducted. This inference is also supported by the nature of the lower crust seismically imaged beneath the Lausitz (Behr et al., 1994). Throughout Saxo-Thuringia, the earlier stages of the Cadomian development are known mainly from drill cores. The Rothstein Formation of Buschmann et al. (1995; Fig. 1) partly makes up the upper continuation of the Lausitz Group. This unit allows two magmatic stages to be distinguished: The first produced enriched MORB (E-MORB) type volcanism in a rift environment that caused hydrothermal alteration of the sediments. Based on an extensional environment, the composition of the detrital input, and widespread slump structures in the older turbiditic graywackes, a back-arc basin setting near a dissective continental magmatic arc was proposed (Buschmann et al., 1998). The second records calcalkaline arc volcanism documented by tuff layers intercalated in
distal basin sediments. Zircon-SHRIMP dating of synsedimentary tuff layers (Buschmann et al., 2001) yielded almost identical ages as those obtained from the Lausitz Group (566 ± 10, 564 ± 3, and 574 ± 8 Ma; Gehmlich et al., 1997). Thus, the Rothstein Formation either represents an earlier, independent rift basin or an early stage of extension and subsidence of the marine basin in which the Lausitz Group was later deposited. The covered Neoproterozoic graywackes between the exposed Lausitz Group and the Rothstein Formation are thought to represent the continuing basin sedimentation of the Lausitz Group, based on closely comparable lithoclast composition (Brause, 1969). Detrital composition and heavy mineral content of the stratigraphically equivalent Leipzig Group (Fig. 1) also indicate a large, more or less uniform source area comprising a dissected magmatic arc (Fig. 12). Sehm (1976) described lithoclasts derived from granitoids, minor volcanic rocks, low-grade metamorphic rocks, and sedimentary rocks, including chert and
134
TABLE 11. SAMPLE LIST FOR GEOCHEMICAL AND DETRITAL COMPONENT ANALYSES Kemnitz Sample 612 639 640 642 643 644 647b 647c 649 653 659 661b 662 667 668 670 678 684 686 692 694 695 698 702 704 709 710 711 712 713 714 715 716 718 719 720 721 722 723 724 725 726 727 729 732 734 735 736 737 738 739 740 741 742 743 744 745 746 747 748 749 750 751 752
Rock type
Sample location
Metagraywacke Metagraywacke Metasiltstone Metatectic hornfels Metabasic rock Metagraywacke Graywacke Graywacke Graywacke Graywacke Metasiltstone Graywacke Cordierite hornfels Cordierite hornfels Metabasic rock Metatectic hornfels Graywacke Graywacke Graywacke Cordierite hornfels Metabasic rock Metagraywacke Metagraywacke Cordierite hornfels Metabasic rock Biotite hornfels Metagraywacke Cordierite hornfels Metabasic rock Anatexite Cordierite hornfels Metatectic hornfels Cordierite hornfels Cordierite hornfels Granodiorite Metagraywacke Porphyric granodiorite Anatexite Porphyric granodiorite Porphyric granodiorite Granodiorit Porphyric granodiorite Anatexite Granodiorite Granodiorite Porphyric granodiorite Granodiorite Granodiorite Porphyric granodiorite Granodiorite Granodiorite Porphyric granodiorite Siltstone Graywacke Graywacke Siltstone Metabasic rock Graywacke Metasiltstone Graywacke Graywacke Siltstone Graywacke Granodiorite
Railway station Arnsdorf Quarry Kindisch Quarry Kindisch Hutberg, Seifersdorf Tobiasmühle, Lotzdorf Tobiasmühle, Lotzdorf Brandhübel, Möhrsdorf Brandhübel, Möhrsdorf Eulenstein, Möhrsdorf Brandhübel, Hennersdorf Brandhübel, Hennersdorf Brandhübel, Hennersdorf Quarry Wiesa, Kamenz Hüttermühle, Radeberg Hüttermühle, Radeberg Castle Klippenburg, Radeberg Heiliger Berg, Hennersdorf Heiliger Berg, Hennersdorf Golksberg, Hennersdorf Wallroda Railway viaduct, Kleinwolmsdorf Railway viaduct, Kleinwolmsdorf Kleinliegau Grosse Röder valley, Schönborn Grosse Röder valley, Liegau Tanneberg, Rammenau Galgenberg, Burkau Galgenberg, Burkau Galgenberg, Burkau Galgenberg, Burkau Burkauer Berg, Burkau Burkauer Berg, Burkau Säuritz Leipsberg, Gödlau Kesselberg, Oberrammenau Oberrammenau Hauffensberg, Hauswalde Hauffensberg, Hauswalde Hohberg, Bretnig Großröhrsdorf Großröhrsdorf Schleißberg, Ohorn Schleißberg, Ohorn Hirschberg, Ohorn Hirschberg, Ohorn Luchsenburg, Ohorn Luchsenburg, Ohorn Hochstein, Pulsnitz Burkauer Berg, Burkau Burkauer Berg, Burkau Napoleonstein, Bischofswerda Goldbach Wohlaer Berg, Wohla Eulenstein, Wohla Wohlaer Berg, Wohla Golksberg, Hennersdorf Golksberg, Hennersdorf Golksberg, Hennersdorf Golksberg, Hennersdorf Lückersdorf Vogelberg, Kamenz Vogelberg, Kamenz Vogelberg, Kamenz Petershain
Coordinates r5428488 r5438320 r5438300 r5421700 r5423560 r5423561 r5436540 r5436540 r5436760 r5436340 r5436790 r5436915 r5438425 r5426085 r5425720 r5425400 r5435625 r5434960 r5435700 r5435765 r5426432 r5426434 r5422860 r5435762 r5422180 r5439380 r5440560 r5440490 r5440561 r5440500 r5441024 r5441070 r5441188 r5441605 r5437892 r5437849 r5436280 r5436300 r5435154 r5431100 r5433000 r5434260 r 5434250 r5434680 r5434565 r5436790 r5436725 r5437708 r5441475 r5441485 r5441538 r5440865 r5437900 r5437625 r5437995 r5436075 r5436085 r5436080 r5435660 r5435383 r5435325 r5434912 r5434911 r5431525
h5661905 h5673880 h5673845 h5669152 h5665820 h5665823 h5636540 h5636541 h5678845 h5676610 h5673590 h5676500 h5681040 h5665695 h5665772 h5665528 h5677935 h5678499 h5679500 h5665205 h5663875 h5663875 h5667640 h5668455 h5668455 h5670578 h5671820 h5671285 h5671819 h5671180 h5670420 h5670403 h5673080 h5674330 h5670110 h5670013 h5670109 h5670090 h5668000 h5666950 h5667008 h5672480 h5672445 h5672730 h5672630 h5671940 h5671965 h5672440 h5669980 h5669985 h5666195 h5667085 h5678440 h5678340 h5678355 h5679150 h5679130 h5679110 h5679520 h5681320 h5682900 h5683190 h5683190 h5682610 Continued
TABLE 11. SAMPLE LIST GEOCHEMICAL AND DETRITAL COMPONENT TheFOR Lausitz graywackes, Saxo-Thuringia, GermanyANALYSES (continued) Sample 753 754 755 756 757 758 759 760 761 762 763 764 765 768 769 771 775 776 777a 777c 778 781 782c 788 789 791 792 793 794 795a 795b 797 798 799 801 802 804 805a 808 809 810 813 814 815 817 818b 819 821 822 823 824 828 829 831 833 834 836 837 839 842 843 844 845 848
Rock type
Sample location
Cordierite hornfels Granodiorite Graywacke Granodiorite Migmatized hornfels Migmatized hornfels Metagraywacke Metatectic hornfels Biotite hornfels Granodiorite Metagraywacke Biotite hornfels Metabasic rock Siltstone Graywacke Metabasic rock Granodiorite Graywacke Graywacke Graywacke Graywacke Anatexite Graywacke Metabasic rock Metatectic hornfels Anatexite Metagraywacke Cordierite hornfels Metagraywacke Cordierite hornfels Anatexite Anatexite Cordierite hornfels Metabasic rock Metatectic hornfels Porphyric granodiorite Anatexite Metabasic rock Metabasic rock Granodiorite Metabasic rock Granodiorite Granodiorite Cordierite hornfels Porphyric granodiorite Metabasic rock Granodiorite Metagraywacke Granodiorite Anatexite Metagraywacke Metabasic rock Granodiorite Metabasic rock Anatexite Metabasic rock Granodiorite Metabasic rock Porphyric granodiorite Migmatized hornfels Biotite hornfels Granodiorite Metabasic rock Metabasic rock
Spitzberg, Schwosdorf Spitzberg, Schwosdorf Bischheim-Häslich Mühlberg, Bischheim-Häslich Drachenberg, Leppersdorf Drachenberg, Leppersdorf Lotzdorf Marienmühle, Seifersdorf Marienmühle, Seifersdorf Marienmühle, Seifersdorf Kleinwolmsdorf Kleinwolmsdorf Kleinwolmsdorf Quarry Butterberg, Bernbruch Quarry Butterberg, Bernbruch Quarry Oßling Spitzberg, Petershain Spitzberg, Petershain Spitzberg, Petershain Spitzberg, Petershain Spitzberg, Petershain Karschberg, Höckendorf Quarry Oßling Quarry Oßling Dresdener Heide Dresdener Heide Schwosdorf Liegau Liegau Liegau-Augustusbad Liegau-Augustusbad Kleinliegau Kleinliegau Grundmühle, Kleinliegau Papiermühle, Seifersdorf Kunathmühle, Seifersdorf Wallroda Wallroda Leppersdorf Quarry Kindisch Quarry Kindisch Rehnsdorf Rehnsdorf Rehnsdorf Schlungwitz, Spree valley Czorneboh, summit Czorneboh, summit Quarry Schwarzkollm Grosse Röder valley, Schönborn Grosse Röder valley, Schönborn Schwosdorf Forkersberg, Rückersdorf Hohwald Pohlaer Berg Sonnenberg, Berge Sonnenberg, Berge Sonnenberg, Berge Mönchswalde Teufelskanzel, Sora Teufelskanzel, Sora Adlerwald, Sora Adlerwald, Sora Adlerwald, Sora Burkauer Berg, Burkau
Coordinates r5431130 r5431160 r5432230 r5431089 r5429273 r5429030 r5423475 r5421924 r5422102 r5422020 r5426932 r5427673 r5427670 r5435213 r5435215 r5423345 r5431180 r5431770 r5431840 r5431841 r5431870 r5425680 r5423350 r5423350 r5418478 r5418455 r5432480 r5423130 r5423127 r5423621 r5423622 r5423100 r5423105 r5422960 r5421330 r5420365 r5428500 r5428502 r5426065 r5438375 r5438260 r5437040 r5437045 r5437073
h5682450 h5682400 h5679700 h5678700 h5670068 h5670411 h5666802 h5668791 h5668923 h5668802 h5663187 h5663121 h5663122 h5685028 h5685026 h5665595 h5682730 h5682999 h5682999 h5682998 h5683025 h5676800 h5665611 h5665612 h5663435 h5663425 h5682235 h5667268 h5667270 h5667274 h5667670 h5667844 h5667842 h5668122 h5669755 h5669944 h5665287 h5665290 h5665033 h5673750 h5673890 h5674950 h5674975 h5674985
r5422763 r5420400 r5432502 r5442800 r5449839 r5444972 r5459918 r5459917 r5459818 r5459252 r5457315 r5457253 r5456907 r5456907 r5456773 r5441400
h5668180 h5669763 h5682266 h5656495 h5657920 h5670688 h5664136 h5664135 h5664071 h5665323 h5665250 h5664937 h5665003 h5665002 h5664808 h5670400 Continued
135
136
TABLE 11. SAMPLE LIST FOR GEOCHEMICALKemnitz AND DETRITAL COMPONENT ANALYSES (continued) Sample
Rock type
Sample location
Coordinates
851 Anatexite Quarry Sora r5457005 852 Porphyric granodiorite Quarry Sora r5457028 853 Metabasic rock Quarry Sora r5457018 867 Siltstone Quarry Oßling r5423348 868a Graywacke Quarry Oßling r5423348 871 Siltstone Quarry Oßling r5423352 877a Graywacke Kunnersdorf-Siebenhufen r5495210 877b Graywacke Kunnersdorf-Siebenhufen r5495211 878 Graywacke Kunnersdorf-Siebenhufen r5495150 886a Graywacke Görlitz r5499228 904 Graywacke Wüsteberg, Gelenau r5433575 905 Graywacke Wahlberg, Gelenau r5433560 907 Graywacke Quarry Oßling r5423350 913a Graywacke Görlitz-Königshufen r5499180 915 Graywacke Görlitz-Königshufen r5499655 916 Graywacke/siltstone Görlitz-Königshufen r5499340 917 Graywacke Görlitz-Königshufen r5499425 918a Graywacke Görlitz-Klingewalde r5498325 919b Graywacke Görlitz-Klingewalde r5498085 921b Graywacke Görlitz-Klingewalde r5498315 932a Metagraywacke Kunnersdorf-Liebstein r5494045 933 Graywacke Kunnersdorf-Liebstein r5493980 934a Graywacke Kapellenberg, Ebersbach r5494940 934b Graywacke Kapellenberg, Ebersbach r5495210 945 Graywacke Kunnersdorf-Siebenhufen r5495080 951 Metasiltstone Prospecting material Wüsteberg r5433570 952 Metasiltstone Prospecting material Wüsteberg r5433570 953 Metagraywacke Prospecting material Wüsteberg r5433570 954 Metasiltstone Prospecting material Wüsteberg r5433570 955 Metagraywacke Prospecting material Wüsteberg r5433570 956 Metabasic rock Prospecting material Wüsteberg r5433570 961 Graywacke Quarry Eiskellerberg, Kalkreuth 962 Graywacke Quarry Eiskellerberg, Kalkreuth 965 Graywacke Quarry Wetterberg, Radeburg r5405505 966 Graywacke Quarry Wetterberg, Radeburg r5405505 967 Graywacke Quarry Butterberg, Bernbruch r5435215 Cgl Conglomerate, enclave Quarry Kindisch r5438375 973 Graywacke Quarry Dobrˇis 975 Graywacke Lhota 976 Graywacke Quarry Libuš, Prague 977 Graywacke Quarry Großzschocher, Leipzig 978 Graywacke Kreuzberg, Röhrsdorf S7 Graywacke Railway station Zgorzelec-Miasto S8 Graywacke Railway station Zgorzelec-Miasto S 11 Metagraywacke Quarry Ujazd S 12 Metagraywacke Quarry Ujazd S 24 Graywacke Railway station Je˛drzychowice 7/22 Metagraywacke Neukirch-Königsbrück 7/45 Graywacke Weißbach-Königsbrück 4755.348 Graywacke Prospecting material Kodersdorf 4755.349 Graywacke Prospecting material Kodersdorf Sb 22 Metagraywacke Weissenberg Sb323a Metagraywacke Großsaubernitz Na2 Metabasic rock Vogelberg, Kamenz Na 3 Metabasic rock Vogelberg, Kamenz Lö1.1 Metabasic rock Stolpen surroundings Lö1.2 Metabasic rock Stolpen surroundings Lö2.1 Metabasic rock Stolpen surroundings Lö2.2 Metabasic rock Stolpen surroundings Lö3.1 Metabasic rock Stolpen surroundings Lö3.2 Metabasic rock Stolpen surroundings Lö3.3 Metabasic rock Stolpen surroundings Lö4.1 Metabasic rock Stolpen surroundings Lö4.2 Metabasic rock Stolpen surroundings Note: r/h—Gauß-Krüger coordinates based on topographic maps 1:10,000 and 1:25,000.
h5665528 h5665588 h5665588 h5665607 h5665608 h5665610 h5673260 h5673260 h5673250 h5669750 h5680660 h5680820 h5665600 h5670495 h5670125 h5669920 h5670690 h5672270 h5672475 h5673040 h5673940 h5674550 h5673100 h5673259 h5673430 h5680663 h5680663 h5680663 h5680663 h5680663 h5680663
h5682008 h5682008 h5685026 h5673750
The Lausitz graywackes, Saxo-Thuringia, Germany
137
QM Lausitz Group
Katzhütte Group Šteˇchovice Group
QM
ck blo al- e ent anc tin ven con pro
plu pro toni ven c-a an rc ce
North-Saxony (Leipzig) Group
tin pro enta ven l-bl an ock ce
vo pro lcani ven c-a an rc ce
QP
con
P
r e c y c le
K
do r og
en
plu
ton
ma
F
ic
gm
ati
vo
c a lcanic rc
Lt Li
magmatic arc
Ls
Figure 12. Figures comparing different Saxo-Thuringian and Teplá-Barrandian Neoproterozoic units to the tectonic setting of their source areas. Ternary systems from Dickinson and Suczek (1979). Abbreviations as in Figure 6.
sandstone fragments (Table 3). He inferred relatively short transport distances and high subsidence and/or erosion rates, which would be typical of an active arc-related tectonic environment. Facies features correspond to a medial fan position. Exposed within the Elbe Zone, the Weesenstein and Clanzschwitz groups (Fig. 1) are fragmentary Cadomian units that became decoupled during the Variscan collision (Linnemann and Schauer, 1999). They represent proximal, upper fan deposits (Linnemann, 1992) and contain conglomeratic horizons that are, in their lower parts, characterized by predominantly granitoids and acid to intermediate volcanic pebbles and a lack of cherts (Schmidt, 1960; Linnemann, 1992). Like the Lausitz conglomeratic enclave described above, these lower parts may represent an older basin stage. The geochemical pattern indicates a continental island-arc provenance (Linnemann et al., 2000). The Thuringian Neoproterozoic (Katzhütte Group; Bankwitz and Bankwitz, 1995) is exposed in a complex of Variscan tectonically stacked Proterozoic and Palaeozoic rocks (Schäfer, 1997; Heuse et al., 2001; Sommer and Katzung, 2004) across the width of the Schwarzburg Structure (“Anticline”; Fig. 1). Formation temperatures are still preserved in detrital white micas of an adequate grain size (>70 μm) despite a greenschist-facies Variscan overprint (Schäfer et al., 2000), and a granitoid derivation could be verified (Kemnitz et al., 1999). In addition to plutonic rock
fragments, the Katzhütte graywackes contain >10% acid to intermediate volcanic rocks and a minor part of chert fragments. This composition not only proves their derivation from a dissected magmatic arc (Fig. 12) but suggests that they were deposited at more or less the same time as the Lausitz Group. Thus, based on geochemical discrimination and isotopic dating, a common tectonic background can be assigned to all of the Neoproterozoic graywacke turbidites from Saxo-Thurinigia, including those parts of the HT–high pressure (HP) Erzgebirge Neoproterozoic units (Mingram, 1998; Tichomirowa et al., 2001; Mingram et al., 2004). Like the Lausitz area, the geochronological data for magmatic rocks and their graywacke host rocks give the same range of Neoproterozoic sedimentation and postcollisional intrusion ages, allowing for a well-constrained stratigraphic correlation (Linnemann et al., 2000). In contrast, the Teplá-Barrandian in the neighboring Czech Republic (Fig. 1) was apparently not part of a common basin with the Saxo-Thuringian units of north Armorica, as previously suspected (Kemnitz, 1994). Although a similar tectonic setting marginal to Proto-Gondwana (Buschmann and Schneider, 1994) and a common parentage from west Africa (Zulauf et al., 1999; Linnemann and Romer, 2002) can be assumed, and furthermore, the Šteˇ chovice graywackes have a remarkably similar detrital composition, a number of crucial
138
Kemnitz
data preclude their correlation. These include: (1) a divergent paleogeographic position (Nance and Murphy, 1994; Murphy et al., 2004), (2) different signatures of the Cadomian magmatism, (3) minor intrusions, (4) a different type of Cadomian metamorphism that lacked large-scale anatectic processes, and (5) different Cambrian basin development (Buschmann and Schneider, 1994; Zulauf et al., 1997). Instead, these features link the Teplá-Barrandian Neoproterozoic successions to South-Armorica (Murphy et al., 2004). CONCLUSIONS The facies record of the Lausitz graywackes indicates a basin setting of continuous extension and deepening, with slight progradation from medial to distal fan position. Hence, the graywackes preserve only part of the Cadomian record. A well-constrained geochronology of Cadomian subduction- and collision-related processes allows this time interval to be bracketed between 560 and 545 Ma (c.f. Linnemann et al., 2004). Supported by geochemical patterns, a more detailed understanding of the depositional setting can be gained through the study of the detritus as well as the mineralogical and structural features of the graywackes. The detrital composition and geochemistry of the psammitic graywacke intervals reflect a dissected magmatic arc provenance and probably a back-arc basin position, with a lithoclast association largely dominated by plutonic rocks. Some granitoidic fragments show a Ca-rich granitic-granodioritic composition and are thought to be derived from pre-Cadomian basement. Another important group of lithoclasts comprises acid to intermediate volcanic rocks and are thought to represent the slightly older members of a bimodal Cadomian arc magmatic suite. An older stratigraphic position for this suite is suggested by a conglomeratic enclave clearly derived from deeper stratigraphic horizons and which is dominated by pebbles of acid to intermediate volcanic rocks. In equivalent Cadomian graywackes of the Teplá-Barrandian, a rhyolitic boulder yielded an age of 570 Ma (Dörr et al., 1992). From the Lausitz region, Tichomirowa et al. (2001) obtained ages between 585 and 565 Ma for mainly acid arc magmatic activity. Members of a basaltic-andesitic suite survive only as enclaves and rafts at an anatectic crustal level that was later exhumed. Within the Lausitz Group, probable products of this basic arc magmatism are represented by tuffaceous graywackes, which are interpreted to represent reworked diluted volcanogenic mass-flow material. Zircon dating (Gehmlich et al., 1997; Buschmann et al., 2001, Linnemann et al., 2004) indicates basaltic (to andesitic) arc volcanic activity at ca 575–560 Ma. In contrast to the Cadomian evolution in the Teplá-Barrandian, the early (rifting) and late (syncollisional) basin stages are not exposed in the Lausitz area. Within the Saxo-Thuringian terrane, a presumably early basin stage is known only from the Rothstein Formation (Buschmann et al., 1995, 2001). In the Lausitz area, however, older, possibly equivalent basin units were wholly involved in the process of anatectic melt generation, which started synchronously with or subsequent to the stage of
arc-magmatic activity (Tichomirowa et al., 2001). As this partial anatexis affected and mixed a large volume of crustal material through mobilization and differentiated intrusions, use of lithostratigraphic nomenclature seems inappropriate. The Neoproterozoic successions of Saxo-Thuringia represent limited time intervals and, hence, fragmentary stages of Cadomian evolution. They are therefore appropriately represented by the recently applied nomenclature, which assigns them to regional groups (Katzhütte Group, etc.). The Lausitz Group represents a tectonostratigraphically uniform, middle to later section of the subduction-related basin stage. To emphasize the common setting of these fragmentary successions and their obvious deposition on the same convergent-margin, it is proposed that they be grouped under a higher ranking term Saxo-Thuringian Supergroup. ACKNOWLEDGMENTS I gratefully acknowledge financial and technical support of the GeoForschungsZentrum (GFZ) Potsdam. I thank Jürgen Eidam (formerly Ernst-Moritz-Arndt University Greifswald, now retired), Ulf Linnemann (Museum für Mineralogie und Geologie Dresden), Bernd Buschmann, and Marion Tichomirowa (Bergakademie Freiberg) for the stimulating discussions we had on field trips and workshops. Thanks are due to Michael Bau (International University Bremen) and Erika Kramer (formerly GFZ Potsdam, now retired) who carried out the REE analyses. The careful reviews and comments of Andrzej Z˙ elaz´niewicz and Damian Nance helped considerably to improve the article. Last but not least I thank Gertraud Budzinski, who has been a valuable field companion and assistant in petrographical work for some years, and the late Gusti Burmann, who always found thoughtful, helpful, and encouraging comments. REFERENCES CITED Bankwitz, P., and Bankwitz, E., 1995, Proterozoikum/Schwarzburger Antiklinorium, in Seidel, G., ed., Geologie von Thüringen: Stuttgart, E. Schweizerbart’sche Verlagsbuchhandlung, p. 46–77. Bea, F., 1980, Geochemistry of biotites in an assimilation process. An approach to recognition of metamorphic biotites from magmatic occurrence: Krystalinikum, v. 15, p. 103–124. Behr, H.-J., Dürbaum, H.-J., Bankwitz, P., Bankwitz, E., Benek, R., Berger, H.-J., et al., eds., 1994, Crustal structure of the Saxothuringian Zone, results of the deep seismic profile MVE-90 (EAST), in DEKORP Research Group (B), The deep reflection seismic profiles DEKORP 3/MVE-90: Zeitschrift für geologische Wissenschaften, v. 22, no. 6, p. 647–669. Behrmann, J.H., 1984, A study of white mica microstructure and microchemistry in a low grade mylonite: Journal of Structural Geology, v. 6, no. 3, p. 283–292, doi: 10.1016/0191-8141(84)90052-X. Bernadová, E., and Cháb, J., 1968, Präassyntische kristalline Schiefer als klastisches Material in jungoberproterozoischen Grauwacken im NW-Teil des Barrandiums: Geologie, v. 17, no. 6/7, p. 753–775. Bhatia, M.R., 1983, Plate tectonics and geochemical composition of sandstone: Journal of Geology, v. 91, p. 611–627. Bhatia, M.R., and Crook, K.A.W., 1986, Trace element characteristics of graywackes and tectonic setting discrimination of sedimentary basins: Contributions to Mineralogy and Petrology, v. 92, p. 181–193, doi: 10.1007/ BF00375292. Bhatia, M.R., and Taylor, S.R., 1981, Trace element geochemistry and sedimentary provinces, a study from the Tasman Geosyncline, Australia: Chemical Geology, v. 33, p. 115–126, doi: 10.1016/0009-2541(81)90089-9.
The Lausitz graywackes, Saxo-Thuringia, Germany Bielicki, K.-H., Haase, G., Eidam, J., Kemnitz, H., Schust, F., Gerstenberger, H., Hiller, H., and Habedank, M., 1989, Pb-Pb and Rb-Sr dating of granitoids from the Lausitz Block, in Fifth working meeting. Isotopes in Nature: Proceedings I: Mitteilungen des Zentralinstituts für Isotopen und Strahlenforschung Leipzig, p. 23–45. Borkowska, M., 1994, Geochemische Charakteristik der Ostlausitzer Granitoide und ihrer Gneisäquivalente in den Sudenten: Terra Nostra, v. 2/94, p. 25–28. Borkowska, M., Hameurt, J., and Vidal, P., 1980, Origin and age of Izera gneisses and Rumburk granites in the Western Sudetes: Acta Geologica Polonica, v. 30, no. 2, p. 121–146. Bouma, A.H., 1962, Sedimentology of some flysch deposits: Amsterdam, Elsevier, 168 p. Brause, H., 1969, Das verdeckte Altpaläozoikum der Lausitz und seine regionale Stellung. Abhandlungen der deutschen Akademie der Wissenschaften Berlin, Klasse Bergbau, Hüttenwesen und Montangeologie, Berlin, v. 1968, no. 1, 143 p. Brigatti, M.F., and Gregnanin, A., 1987, Crystal chemistry of igneous rock biotites: Mineralogy and Petrology, v. 37, p. 323–341, doi: 10.1007/BF01161824. Burmann, G., 1966, Mikroreste aus der Lausitzer Grauwackenformation: Monatsberichte der deutschen Akademie der Wissenschaften Berlin, v. 8, no. 3, p. 218–224. Burmann, G., 1969, Organische Mikrofossilien in präkambrischen Sedimenten Sachsens und Thüringens: Monatsberichte der deutschen Akademie der Wissenschaften Berlin, v. 11, no. 4, p. 297–307. Burmann, G., 1972, Problematika aus der Lausitzer Grauwackenformation: Jahrbuch für Geologie, v. 4, p. 387–423. Burmann, G., 1997, Präkambrisches Algenbenthos aus Kieselpeliten. Teil 1, Palisadia wagneri n. gen. n. sp.: Zeitschrift für geologische Wissenschaften, v. 25, no. 1/2, p. 41–89. Burmann, G., 2000, Die biostratigraphische Gliederung der präkambrischen Lausitzer Grauwacken-Einheit. Teil 1. Das Lausitzer Antiklinorium und Nachbargebiete (nordsächsisches Antiklinorium, Jena–Pegau, Schwarzburger Antiklinorium): Zeitschrift für geologische Wissenschaften, v. 28, no. 1/2, p. 33–69. Burmann, G., 2001, Die biostratigraphische Gliederung der präkambrischen Lausitzer Grauwacken-Einheit. Teil 2, Zur überregionalen Korrelation (Osteuropäische Tafer, Teplá-Barrandium, Bretagne-Normandie, Skandinavien): Zeitschrift für geologische Wissenschaften, v. 29, no. 3, p. 299–333. Buschmann, B., and Schneider, J., 1994, Gedanken zu cadomischen Terranen der Böhmischen Masse, in Abstracts, DFG-Schwerpunktprogramm “Orogene Prozesse—ihre Quantifizierung und Simulation am Beispiel der Varisciden,” 29–31 March, Würzburg: Terra Nostra, v. 94, no. 5, p. 31–32. Buschmann, B., Linnemann, U., Schneider, J., and Süss, T., 1995, Die cadomische Entwicklung im Untergrund der Torgau-Doberluger Synklinale: Zeitschrift für geologische Wissenschaften, v. 23, no. 5/6, p. 729–749. Buschmann, B., Jonas, P., Herzig, P., and Schneider, J., 1998, Geotectonic facies analysis of the Rothstein Formation, a fragment of a Cadomian continental margin backarc basin: Schriften des Staatlichen Museums für Mineralogie und Geologie, Dresden, v. 9, p. 109–111. Buschmann, B., Nasdala, L., Jonas, P., Linnemann, U., and Gehmlich, M., 2001, SHRIMP U-Pb dating of tuff-derived and detrital zircons from Cadomian marginal basin fragments (Neoproterozoic) in the northeastern Saxothuringian Zone (Germany): Neues Jahrbuch für Geologie und Paläontologie Monatshefte, v. 6, p. 321–342. Chauvel, J.J., and Mansúy, C., 1981, Micropaléontologie du Protérozoique du Massif Armoricain (France): Precambrian Research, v. 15, no. 1, p. 25– 42, doi: 10.1016/0301-9268(81)90065-6. Chlupácˇ, I., 1993, Geology of the Barrandian: Frankfurt am Main, Senckenberg-Buch, v. 69, p. 163. Dickinson, W.R., 1970, Interpreting detrital modes of graywackes and arkose: Journal of Sedimentary Petrology, v. 40, no. 2, p. 695–707. Dickinson, W.R., and Suczek, C.A., 1979, Plate tectonics and sandstone compositions: American Association of Petroleum Geologists Bulletin, v. 63, p. 2164–2187. Dickinson, W.R., and Valloni, R., 1980, Plate tectonic settings and provenance of sands in modern ocean basins: Geology, v. 8, no. 2, p. 82–86, doi: 10.1130/0091-7613(1980)8<82:PSAPOS>2.0.CO;2. Dickinson, W.R., Ingersoll, R.V., Cowan, D.S., Helmold, K.P., and Suczek, C.A., 1982, Provenance of Franciscan graywackes in coastal California: Geological Society of America Bulletin, v. 93, p. 95–107. Dörr, W., Fišera, M., and Franke, W., 1992, Cadomian magmatic events in the Bohemian Massif—U-Pb data from felsic magmatic pebbles, in Abstracts
139
7th Geological Workshop “Styles of superposed Variscan nappe tectonics”: Prague, Czechoslovakian Academy of Science, p. 3. Ebert, H., 1943, Das granitische Grundgebirge der östlichen Lausitz: Preisschriften, gekrönt und herausgegeben von der Fürstlich Jablonowskischen Gesellschaft zu Leipzig, v. 58, 119 p. Eidam, J., 1988, Zur anatektischen und magmatischen Entwicklung der granitoiden Gesteine der Lausitzer Antiklinalzone: Zeitschrift für geologische Wissenschaften, v. 16, p. 393–406. Eidam, J., Hammer, J., Korich, D., and Bielicki, K.-H., 1995, Amphibolebearing granites in the Lausitz Anticlinal Zone, Variscan I-type magmatism at the northern margin of the Bohemian Massif: Neues Jahrbuch für Mineralogische Abhandlungen, v. 168, p. 259–281. Fiala, F., 1948, Conglomérats algonkiens de la Bohème central: Prag, Sbornik Stát geologický ústav CˇSR, v. 15, p. 612–675. Foster, M.D., 1960, Interpretation of the composition of trioctahedral micas: Washington, D.C., U.S. Geological Survey Professional Paper 354-B, p. 11–49. Füchtbauer, H., 1988, Sedimente und Sedimentgesteine: Stuttgart, E. Schweizerbart’sche Verlagsbuchhandlung, 1141 p. Gehmlich, M., Tichomirowa, M., Linnemann, U., Lützner, H, and Bombach, K., 1997, Kalibrierung und Korrelation des saxothuringischen Ordoviziums durch die Einzelzirkon-Evaporationsmethode, in Abstracts, SPPKolloquium Orogene Prozesse, ihre Quantifizierung und Simulation am Beispiel der Varisciden, 13–14 March, Bayreuth: Terra Nostra, v. 97, no. 5, p. 43–46. Guidotti, C.V., and Sassi, F.P., 1976, Muscovite as a petrogenetic indicator mineral in pelitic schists: Neues Jahrbuch Mineralogische Abhandlungen, v. 127, p. 97–142. Guidotti, C.V., and Sassi, F.P., 2002, Constraints on studies of metamorphic KNa white micas, in Micas, crystal chemistry and metamorphic petrology: Reviews in Mineralogy and Geochemistry, v. 46, p. 413–448. Hammer, J., 1996, Geochemie und Petrogenese der cadomischen und spätvariszischen Granitoide der Lausitz: Freiberger Forschungsheft, v. C463, p. 1–107. Hammer, J., Eidam, J., Krauss, M., and Korich, D., 1998, The Lausitz (northern margin of the Bohemian Massif) as part of the “Avalonian-Cadomian Belt”: Schriften des Staatlichen Museums für Mineralogie und Geologie Dresden, v. 9, p. 143. Hammer, J., Eidam, J., Röber, B., and Ehling, B.-C., 1999, Prävariscischer granitoider Magmatismus am NE-Rand des Böhmischen Massivs—Geochemie und Petrogenese: Zeitschrift für geologische Wissenschaften, v. 27, no. 5/6, p. 401–415. Heuse, T., Kroner, U., and Bartl, S., 2001, Tektono- und lithostratigraphische Kartierungskriterien für den Zentralteil des Schwarzburger Sattels: Geowissenschaftliche Mitteilungen Thüringen, v. 9, p. 107–124. Hirschmann, G., 1966, Assyntische und variszische Baueinheiten im Grundgebirge der Oberlausitz (unter spezieller Berücksichtigung der Geologie des östlichen Görlitzer Schiefergebirges): Freiberger Forschungsheft, v. C212, p. 1–146. Ingersoll, R.V., and Suczek, C.A., 1979, Petrology and provenance of Neogene sand from Nicobar and Bengal fans, DSDP sites 211 and 218: Journal of Sedimentary Petrology, v. 49, p. 1217–1228. Ingersoll, R.V., Bullard, T.F., Ford, R.L., and Pickle, J.D., 1984, The effect of grain size on detrital modes: A test of the Gazzi-Dickinson point-counting method: Journal of Sedimentary Petrology, v. 54, p. 103–116. Kemnitz, H., 1994, Provenance and plate tectonic interpretation of Upper Brioverian turbidites in southeast Germany (Lausitz Group): Zentralblatt für Geologie und Paläontologie I, v. 1992, p. 803–814. Kemnitz, H., 1997, Lausitzer Masse. Proterozoikum, in Hoth, K., Berger, K.-H., and Mund, G., eds., Stratigraphie von Deutschland II. Ordovizium, Kambrium, Vendium, Riphäikum. Teil I, Thüringen, Sachsen, Ostbayern: Courier Forschungs-Institut Senckenberg, v. 200, p. 323–324. Kemnitz, H., 1998, Some remarks on the Neoproterozoic Lausitz greywackes, Germany, and their mafic intercalations: Acta universitae Carolinae— Geologica, v. 42, no. 3/4, p. 443–446. Kemnitz, H., and Budzinski, G., 1991, Beitrag zur Lithostratigraphie und Genese der Lausitzer Grauwacken: Zeitschrift für geologische Wissenschaften, v. 19, no. 4, p. 433–441. Kemnitz, H., and Budzinski, G., 1994, Die Grauwacken der Lausitz und ihre cadomische Prägung: Abhandlungen des Staatlichen Museums für Mineralogie und Geologie Dresden, v. 40, p. 37–98. Kemnitz, H., Gottesmann, B., and Naumann, R., 1994, Detrital micas from Upper Broverian greywackes of Lusatia—Chemistry and first interpretation: Zentralblatt für Geologie und Paläontologie, Teil 1, v. 1992, p. 837–846.
140
Kemnitz
Kemnitz, H., Schäfer, F., Romer, R.L., and Rosner, M., 1999, Die geodynamische Geschichte eines variszischen Retrokeils im NW-Teil des Saxothuringikums unter besonderer Berücksichtigung des Vesser-Gebietes: Beiträge zur Geologie von Thüringen, Neue Folge, v. 6, p. 41–90. Kemnitz, H., Romer, R.L., and Oncken, O., 2002, Gondwana break-up and the northern margin of the Saxothuringian belt (Variscides of central Europe): International Journal of Earth Sciences, v. 91, no. 2, p. 246–259, doi: 10.1007/s005310100209. Klápová, H., and Hyršl, J., 2000, Correlation of Neoproterozoic conglomerates of the Barrandian, Saxothuringian and Moldanubian units: Veˇstnik Cˇesky geologicky ústav, v. 75, no. 3, p. 217–228. Konzalova, M., 1974, Some new microorganisms from the Bohemian Precambrian (Upper Proterozoic) and Middle Cambrian: Review of Paleobotany and Palynology, v. 18, p. 41–56, doi: 10.1016/0034-6667(74)90007-4. Korytowski, A., Dörr, W., and Z˙elaz´niewicz, A., 1993, U-Pb dating of (meta)granitoids in the NW Sudetes (Poland) and their bearing on tectonostratigraphic correlation: Terra Nova, v. 5/93, p. 331–332. Krauss, M., Eidam, J., Hammer, J., and Korich, D., 1992, Die cadomisch-variszische Entwicklung des Lausitzer Granodiorit-Komplexes: Zentralblatt für Geologie und Paläontologie I, v. 1992, no. 1/2, p. 71–85. Kröner, A., Hegner, E., Hammer, J., Haase, G., Bielicki, K.-H., Krauss, M., and Eidam, J., 1994, Geochronology and Nd-Sr systematics of Lausitz granitoids, significance for the evolution of the Variscan orogen in east-central Europe: Geologische Rundschau, v. 83, no. 2, p. 357–376. LeBas, M.J., LeMaitre, R.W., Streckeisen, A., and Zanettin, B., 1986, Chemical classification of volcanic rocks based on the total alkali-silica diagram: Journal of Petrology, v. 27, p. 745–750. Linnemann, U., 1992, Glazioeustatisch kontrollierte Sedimentationsprozesse im Oberen Proterozoikum der Elbezone (Weesensteiner Gruppe/Sachsen): Zentralblatt für Geologie und Paläontologie I, v. 1991, no. 12, p. 2907–2934. Linnemann, U., and Buschmann, B., 1995, Die cadomische Diskordanz im Saxothuringikum (oberkambrisch–tremadocische overlap-Sequenzen): Zeitschrift für geologische Wissenschaften, v. 23, no. 5/6, p. 707–727. Linnemann, U., and Romer, R.L., 2002, The Cadomian orogeny in SaxoThuringia, Germany, geochemical and Nd-Sr-Pb isotopic characterization of marginal basins with constraints to geotectonic setting and provenance: Tectonophysics, v. 352, p. 33–64, doi: 10.1016/S0040-1951(02)00188-9. Linnemann, U., and Schauer, M., 1999, Die Entstehung der Elbezone vor dem Hintergrund der cadomischen und variszischen Geschichte des Saxothuringischen Terranes—Konsequenzen aus einer abgedeckten geologischen Karte: Zeitschrift für geologische Wissenschaften, v. 27, no. 5/6, p. 529–561. Linnemann, U., Gehmlich, M., Tichomirowa, M., Buschmann, B., Jonas, P., Lützner, H., and Bombach, K., 2000, From Cadomian subduction to Early Palaeozoic rifting, the evolution of Saxo-Thuringia at the margin of Gondwana in the light of single zircon geochronology and basin development (Central European Variscides, Germany), in Franke, W., Haak, V., Oncken, O., and Tanner, D., eds., Orogenic processes, quantification and modelling in the Variscan belt: London, Geological Society of London Special Publication 179, p. 131–153. Linnemann, U., McNaughton, N.J., Romer, R.L., Gehmlich, M., Drost, K., and Tonk, C., 2004, West African provenance for Saxo-Thuringia (Bohemian Massif). Did Armorica ever leave pre-Pangean Gondwana?—U/Pb-SHRIMP zircon evidence and the Nd-isotopic record: International Journal of Earth Sciences, v. 93, no. 5, p. 683–705, doi: 10.1007/s00531-004-0413-8. Lobst, R., 1996, Zum Geröllbestand des Konglomerat-Einschlusses im Granodiorit von Kindisch: Veröffentlichung des Museums der Westlausitz Kamenz, v. 19, p. 19–25. Löffler, H.K., 1982 Die eruptiven und metamorphen Gesteine des Lausitzer Blocks; Teil 2, Evolution und Genese der intrusiven basischen Magmatite vom Valtengrund am Hohwald/Oberlausite: Zeitschrift für Geologische Wissenschaften, v. 10, p. 217–229. Lorenz, W., and Burmann, G., 1972, Alterskriterien für das Präkambrium am Nordrand der Böhmischen Masse. Teil I, Zur Abgrenzung und absoluten Altersdatierung: Geologie, v. 21, no. 4/5, p. 405–417. Massonne, H.J., and Schreyer, W., 1987, Phengite geobarometry based on the limiting assemblage with K-feldspar, phlogopite, and quartz: Contributions to Mineralogy and Petrology, v. 96, p. 212–224, doi: 10.1007/BF00375235. Mazur, S., and Kryza, R., 1996, Superimposed compressional and extensional tectonics in the Karkonosze-Izera block, NE Bohemian Massif, in Oncken, O., and Janssen, C., eds., Basement tectonics 11: Europe and other regions: Dordrecht, Kluwer Academic, p. 51–66.
McLennan, S. M., 2001, Relationship between the trace element composition of sedimentary rocks and upper continental crust: Geochemistry, Geophysics, Geosystems, v. G3, no. 2, p. 1525–2027. Meschede, M., 1986, A method of discriminating between different types of mid-ocean ridge basalts and continental tholeiites with the Nb-Zr-Y diagram: Chemical Geology, v. 56, no. 3–4, p. 207–218. Mingram, B., 1998, The Erzgebirge, Germany, a subducted part of northern Gondwana, geochemical evidence for repetition of early Paleozoic metasedimentary sequences in metamorphic thrust units: Geological Magazine, v. 135, no. 6, p. 785–801, doi: 10.1017/S0016756898001769. Mingram, B., Kröner, A., Hegner, E., and Krentz, O., 2004, Zircon ages, geochemistry, and Nd isotopic systematics of pre-Variscan orthogneisses from the Erzgebirge, Saxony (Germany), and geodynamic interpretation: International Journal of Earth Sciences, v. 93, no. 5, p. 706–727, doi: 10.1007/s00531-004-0414-7. Möbus, G., 1956, Petrographisch-tektonische Untersuchungen im Lausitzer Granitmassiv: Abhandlungen der Deutschen Akademie der Wissenschaften, Klasse Chemie, Geologie und Biologie, v. 1955, p. 8. Murphy, J.B., Pisarevsky, S.A., Nance, R.D., and Keppie, J.D., 2004, Neoproterozoic–Early Paleozoic evolution of peri-Gondwanan terranes, implications for Laurentia-Gondwana connections: International Journal of Earth Sciences, v. 93, no. 5, p. 659–682, doi: 10.1007/s00531-004-0412-9. Nance, R.D., and Murphy, J.B., 1994, Contrasting basement isotopic signatures and the palinspastic restoration of peripheral orogens: Example from the Neoproterozoic Avalonian-Cadomian belt: Geology, v. 22, p. 617–620, doi: 10.1130/0091-7613(1994)022<0617:CBISAT>2.3.CO;2. Nance, R.D., Murphy, J.B., Strachan, R.A., D´Lemos, R.S., and Taylor, G.K., 1991, Late Proterozoic tectonostratigraphic evolution of the Avalonian and Cadomian terranes: Precambrian Research, v. 52, p. 41–78. Nasdala, L., and Pfeiffer, L., 1991, Zu den “Kaulen” aus der Grauwackes des Vogelberges bei Kamenz/Sachsen: Veröffentlichung des Museums der Westlausitz Kamenz, v. 15, p. 3–14. Pearce, J.A., and Cann, J.R., 1973, Tectonic setting of basic volcanic rocks determined using trace element analysis: Earth and Planetary Science Letters, v. 19, p. 290–300, doi: 10.1016/0012-821X(73)90129-5. Rathner, U., 2001, Ergebnisse tektonischer Untersuchungen in der jungproterozoischen Grauwackenfolge des Tagebaus Halbach am Butterberg bei Bernbruch (Westlausitz/Sachsen): Geoprofil, v. 10, p. 12–38. Rimann, E., 1910, Der geologische Bau des Isergebirges und seines nördlichen Vorlandes: Jahrbuch der Preussischen Geologischen Landesanstalt Berlin, v. 31, no. 1, p. 482–533. Röhlich, P., 1964, Submarine slides and mudflows in the youngest Algonkian of Central Bohemia (Czech.): Sbornik Geologickych Veˇ d Geologie, v. 6, p. 89–121. Roser, B.P., and Korsch, R.J., 1986, Determination of tectonic setting of sandstone-mudstone suites using SiO2-content and K2O/Na2O-ratio: Journal of Geology, v. 94, no. 5, p. 635–650. Schäfer, F., 1997, Krustenbilanzierung eines variscischen Retrokeils im Saxothuringikum: Potsdam, Scientific Technical Report GFZ Potsdam 16, 130 p. Schäfer, F., Oncken, O., Kemnitz, H., and Romer, R.L., 2000, Upper-plate deformation during collisional orogeny, a case study from the German Variscides (Saxo-Thuringian Zone), in Franke, W., Haak, V., Oncken, O., and Tanner, D., eds., Orogenic processes, quantification and modelling in the Variscan belt: London, Geological Society of London Special Publication 179, p. 281–302. Schmidt, K., 1960, Die Geröllführung algonkisch-kambrischer Grauwacken des Westlausitzer Zuges: Freiberger Forschungsheft, v. C91, p. 1–98. Schöbel, H., 1985, Schichtflächenmarken und synsedimentäre Deformationsgefüge in der riphäischen Grauwacke der Kamenzer Serie: Veröffentlichung des Museums der Westlausitz Kamenz, v. 9, p. 14–42. Schwab, G., 1962, Klufttektonische Untersuchungen der Nordlausitzer Grauwackenformation unter Berücksichtigung der Gesteinsklüftung des Lausitzer Zweiglimmergranits: Abhandlungen der deutschen Akademie der Wissenschaften Berlin, Klasse Chemie, Geologie, Biologie, v. 1962, no. 2, p. 1–80. Sehm, K., 1972, Lithofazielle und minerogenetische Untersuchungen an der Grauwacken-Pelit-Folge des Nordsächsischen Sattels [Ph.D. diss.]: Freiberg, Bergakademie Freiberg, 269 p. Sehm, K., 1976, Lithologisch-petrofazielle und metallogenetische Untersuchung der Grauwacken-Pelit-Folge des Nordsächsischen Antiklinoriums: Freiberger Forschungsheft. v. C311, p. 5–134. Seston, R., Winchester, J.A., Piasecki, M.A.J., Crowley, Q.G., and Floyd, P.A., 2000, A structural model for the western-central Sudetes, a deformed
The Lausitz graywackes, Saxo-Thuringia, Germany stack of Variscan thrust sheets: Journal of the Geological Society of London, v. 157, p. 1155–1168. Sommer, M., and Katzung, G., 2004, Abgrenzung, Gliederung und strukturelle Entwicklung des Neoproterozoikums im Schwarzburger Antiklinorium (Thüringisch-Fränkisch-Vogtländisches Schiefergebirge, Saxothuringische Zone): Zeitschrift für geologische Wissenschaften, v. 32, no. 2–4, p. 191–226. Stephens, M.B., Glasson, M.J., and Keays, R.R., 1979, Structural and chemical aspects of metamorphic layering development in metasediments from Clunes, Australia: American Journal of Science, v. 279, p. 129–160. Taylor, S.R., and McLennan, S.M., 1985, The continental crust, its composition and evolution: Oxford, Blackwell Scientific Publications, 312 p. Tichomirowa, M., Berger, H.-J., Koch, E.A., Belyatski, B.V., Götze, J., Kempe, U., Nasdala, L., and Schaltegger, U., 2001, Zircon ages of highgrade gneisses in the Eastern Erzgebirge (Central European Variscides)—Constraints on origin of the rocks and Precambrian to Ordovician magmatic events in the Variscan foldbelt: Lithos, v. 56, p. 303–332, doi: 10.1016/S0024-4937(00)00066-9. Velde, B., 1967, Si+4 content of natural phengites: Contributions to Mineralogy and Petrology, v. 14, p. 250–258, doi: 10.1007/BF00376643. Villa, I.M., 1998, Isotopic closure: Terra Nova, v. 10, p. 42–47, doi: 10.1046/ j.1365-3121.1998.00156.x. von Gaertner, H.-R., 1964, Einige Beobachtungen zum Alter des Isergneises: Neues Jahrbuch Geologie und Paläontologie Monatshefte, v. 5, no. 64, p. 257–270. Walker, R.G., 1979, Facies models: Geoscience Canadian Reprint Series v. 1, p. 1–211. Weber, B., Kemnitz, H., and Paech, K., 1990, Zur Kenntnis der Mikrofossilien und Problematika aus der Lausitzer Grauwacke (Kamenzer Gruppe): Veröffentlichungen des Museums der Westlausitz Kamenz, v. 14, p. 9–34.
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Wimmenauer, W., 1984, Exkursionen C3 und C4; Das praevariskische Kristallin im Schwarzwald, in Exkursionsfuehrer; 62. Jahrestagung der Deutschen Mineralogischen Gesellschaft, gemeinsam mit der Oesterreichischen Mineralogischen Gesellschaft und der Schweizerischen Mineralogischen und Petrographischen Gesselschaft: Fortschritte der Mineralogie, Beiheft, v. 62, no. 2, p. 69–86. Wybrecht, E., Duplay, J., Pique, A., and Weber, F., 1985, Mineralogical and chemical evolution of white micas and chlorites, from diagenesis to low-grade metamorphism, data from various size fractions of greywackes (Middle Cambrian, Morocco): Mineralogical Magazine, v. 49, no. 352, p. 401–411. Z˙aba, J., 1985, Progressive regional metamorphism of the Izera Block, Western Sudetes (Poland): Acta Universitatis Carolinae, Geology, v. 1, p. 63–88. Z˙elaz´niewicz, J., Dörr, W., Bylina, P., Franke, W., Haack, U., Heinisch, H., Schastok, J., Grandmontagne, K., and Kulicki, C., 2004, The eastern continuation of the Cadomian orogen, U-Pb zircon evidence from SaxoThuringian granitoids in south-western Poland and the northern Czech Republic: International Journal of Earth Sciences, v. 93, no. 5, p. 773– 781, doi: 10.1007/s00531-004-0418-3. Zulauf, G., Dörr, W., Fiala, J., and Vejnar, Z., 1997, Late Cadomian crustal tilting and Cambrian transtension in the Teplá-Barrandian unit (Bohemian Massif, Central European Variscides): Geologische Rundschau, v. 86, no. 3, p. 571–584, doi: 10.1007/s005310050164. Zulauf, G., et al., 1999, Age constraints on the Cadomian evolution of the TepláBarrandian unit (Bohemian Massif) through electron microprobe dating of metamorphic monazite: ZDGG, v. 150, no. 4, p. 627–639.
MANUSCRIPT ACCEPTED BY THE SOCIETY 3 OCTOBER 2006
Printed in the USA
Geological Society of America Special Paper 423 2007
Paleontological data from the Early Cambrian of Germany and paleobiogeographical implications for the configuration of central Perigondwana Olaf Elicki* Freiberg University, Geological Institute, 09599 Freiberg, Germany
ABSTRACT Fossiliferous Early Cambrian strata from central Europe are known from two fragmentary preserved units situated in eastern Germany: the Görlitz syncline and the Torgau-Doberlug syncline. The fossil assemblages from both regions and their geological framework are presented. The taxonomic content, the internal structure, and the biostratigraphic positions of the containing assemblages are quite different. Trilobites and other shelly fossils (from Görlitz syncline) and archaeocyathans (from Torgau-Doberlug syncline) show strong relationships to equivalent faunas in Morocco, Spain, and France. Further, there are some indications for relationships to the Far East. The German faunas indicate very active faunal exchange within the periGondwanan realm and probably over longer distances along the Gondwana margin, too. Together with the nearly shelf-wide consistent sedimentary facies patterns, consequently, they contradict the model of local and isolated basins for the areas of deposition. The current patchy geographical distribution of Cambrian sediments in central and southern Europe is interpreted as a phenomenon of (1) regionally different sedimentation rates on the Gondwanan shelf (probably by local origination of open intra-shelf basins caused by a general rifting process), and of (2) Late Cambrian to Early Ordovician different-scaled uplift and denudation of parts of the shelf. The content, the coinciding evolutionary patterns, and the paleogeographical relationships of the Cambrian faunas suggest that separation and evolution of terranes in central Perigondwana had not started before the end of the Cambrian or the beginning of the Early Ordovician. For Early and Middle Cambrian times at least, a slightly differentiated shelf-configuration of Perigondwana without isolated areas (terranes) fits best with the paleontological and sedimentological data. Keywords: Cambrian, Perigondwana, paleogeography, Görlitz syncline, TorgauDoberlug syncline
*E-mail:
[email protected]. Elicki, O., 2007, Paleontological data from the Early Cambrian of Germany and paleobiogeographical implications for the configuration of central Perigondwana, 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. 143–152, doi: 10.1130/2007.2423(05). For permission to copy, contact
[email protected]. ©2007 Geological Society of America. All rights reserved.
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INTRODUCTION The geological roots of the European basement areas lie in the Neoproterozoic so-called “Avalonian-Cadomian orogenetic belt,” which was situated at the periphery of Gondwana (Nance and Murphy, 1994; Torsvik et al., 1996). In Cambrian times the central European basement areas indicate first rifting tendency (ultramafic, mafic, and granitoid rocks), which—after widely accepted models—finally led to the break-up of peri-Gondwanan terranes (Linnemann et al., 2000; Kemnitz et al., 2002). Generally, fossiliferous Cambrian strata are very rare in Germany. They are known from the Görlitz and the Leipzig areas, from the Franconian Forest, and from Thuringia (Fig. 1). The Thuringia deposits (two single research drillings from the 1960s) are not considered here because of (1) the very poor preservation of the few microfossil remains found and (2) the lack of biostratigraphic resolution of their general Cambrian age (Blumenstengel, 1980). All these deposits belong to the Saxo-Thuringian terrane (sensu Linnemann and Schauer, 1999). Among them, successions of biostratigraphically indicated Early Cambrian age are known from two regions: from the Görlitz area (Görlitz syncline, a small area near the southeastern border to Poland) and from the Leipzig area (Torgau-Doberlug syncline). These successions were depos-
5°N
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an m r Ge
y
Berlin TDS
GS
Leipzig
TH
Würzburg
Dresden
50°E
FF
Munich
46°E
Figure 1. Physical geographic map of Germany, indicating regions containing Cambrian faunas mentioned in the text. FF—Franconian Forest; GS—Görlitz syncline; TDS—Torgau-Doberlug syncline; TH—location of two drillings in Thuringia providing Cambrian fossils.
ited on or at the flank of the Cadomian consolidated basement (Lusatian block; Linnemann and Buschmann, 1995; Linnemann and Schauer, 1999; Jonas et al., 2000). The structural characteristic of the relations between the Cadomian basement and the marginal synclines is still a matter of discussion, mainly because of the poor outcrop conditions in the area (e.g., Franke, 1984; Buschmann et al., 1995, 2006; Linnemann and Buschmann, 1995; Linnemann and Schauer, 1999; Göthel, 2001). Paleontological studies concerning the German Cambrian faunas have a long tradition. Initially, deeper investigations were focused on the Middle Cambrian successions of the Franconian Forest and the Leipzig area (Torgau-Doberlug syncline) and, later on, on Early Cambrian deposits of the Görlitz and TorgauDoberlug synclines. Compilations of the state of knowledge at different times were given by Sdzuy (1972, for the so-called “acadobaltic province”) and by Elicki (1997). GEOLOGICAL SETTING The poorly exposed succession of the Görlitz syncline (Charlottenhof Formation; Fig. 2) is characterized by a suite of shallow-marine massive dolostones (up to 100 m) and overlying bedded limestones (up to 80 m) called the Ludwigsdorf Member (Elicki and Schneider, 1992; Elicki, 1994). Siliciclastics (90– 120 m) follow to the top (Lusatiops Member). Because of the complex tectonic situation, the stratigraphic continuations above and below are unknown. Because of a strong diagenetic overprint the lower portion of the Ludwigsdorf Member (dolostones) does not yield any biotic remains. Former reports of problematic structures interpreted as probable “archaeocyathans” by Schwarzbach (1934) are rejected after critical re-evaluation (Elicki, 1997). The overlying bedded, and in its highest part nodular, limestones, in contrast, show many sedimentary patterns and are partly rich in fossils. These patterns allow a distinction of two facies realms: (1) shallow and open–marine facies and (2) restricted lagoonal facies (Elicki and Schneider, 1992). The former is characterized by more or less siliciclastic-influenced, bioclastic wackestone, packstone, and floatstone with phosphatic black pebbles, load casts, wave ripples, cross bedding, small channels, and a rich shelly fauna (Elicki and Schneider, 1992; see also below). The latter shows a regular alternation of “Zebra limestone” and dolomitic limestone with slumping and tepee-structures and a primary sulfate content (Elicki and Schneider, 1992). Besides some rare allochthonous remains, this second facies type is free of fossils. The overlaying Lusatiops Member is represented by red and green claystone and siltstone with minor sandstone intercalations. The fossil content is dramatically reduced (some trilobites, brachiopods, and hyoliths). Taphonomic features suggest a rather quiet and deeper shelf environment than those of the underlying Ludwigsdorf Member (Geyer and Elicki, 1995). The Early Cambrian succession of the Torgau-Doberlug syncline (subsurface outcrops; Fig. 3) has a thickness estimated to be ~700–1000 m (Brause, 1969), overlain by Middle Cambrian
???
Delitzsch Formation Tröbitz Formation
claystone
basic volcanics
siltstone
dolostone
sandstone
limestone
local debris flows
nodular limestone
no record
sandy limestone
?
?
?
?
?
?
?
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Rosenfeld Member
Zwethau Formation 100 m
Figure 2. Lithological column of the Cambrian succession of the Görlitz syncline (after Elicki, 2005). Legend valid for Figures 2 and 3. Mbr.—member.
?
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Torgau Member
Lus atiops Mbr. upper Ludwigsdorf Mbr.
lower Ludwigsdorf Mbr.
100 m
Charlottenhof Formation
Paleontological data from the Early Cambrian and paleobiogeographical implications
Figure 3. Lithological column of the Cambrian succession of the Torgau-Doberlug syncline (after Buschmann et al., 2006). See Figure 2 for legend.
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Early Cambrian
strata of probably about the same thickness and covered by up to 200 m of Cenozoic rocks. Locally, Carboniferous sediments overlay the Cambrian strata. Early Cambrian deposition started after the Cadomian unconformity, with locally developed debris flows, followed by shallow-marine carbonates and subordinate siliciclastics (Zwethau Formation). In the lower part of the succession carbonates distinctly predominate (Torgau Member), whereas the upper part is characterized by alternations of siliciclastic-carbonate rocks and pure siliciclastics (Rosenfeld Member). Rarely, intrusive volcanics (diabases) occur (Freyer and Suhr, 1987; Buschmann et al., 1995; Elicki, 1999a,b). The Torgau Member consists of up to 500 m or more of fossiliferous limestones and dolostones. Calcimicrobial, oolitic, and intraclastic limestones are common. Cyanobacteria and archaeocyathans are widespread, and some skeletal fossils were also recovered (Sdzuy, 1962; Freyer and Suhr, 1987, 1992; Elicki, 1992, 1994, 1999b; Elicki and Debrenne, 1993). Shallow depositional conditions are indicated by wave ripples, cross bedding, mudcracks, local evaporites, and biofacies (see below). The environment is interpreted as a shallow subtidal to intertidal ramp with shoals and lagoons (Elicki, 1992, 1999a,b). The poorly investigated overlying Rosenfeld Member (up to 280 m) is represented by an alternation of limestones, dolostones, and siliciclastics (with some diabases in the upper part). Freyer and Suhr (1987) reported a probable higher amount of siliciclastics and of redeposited materials (archaeocyathans,
cyanobacteria, shelly remains) as well as graded carbonates, which led these authors to interpret a transitional facies into a “basinal” environment. Because of the complex structure of the Torgau-Doberlug syncline, the transition into the overlying sandstones and claystones of the Middle Cambrian Tröbitz Formation and Delitzsch Formation remains unclear (Brause, 1969; Elicki, 1997). Regardless an outstanding detailed sedimentological investigation of this siliciclastic succession, a deeper shelf environment below stormwave base is assumed, based on the cyclicity of deposition and of the containing fauna (some trilobites and traces, occasional brachiopods, hyoliths, and others; see Table 2 below). EARLY CAMBRIAN BIOTA The fossils of the Charlottenhof Formation (Görlitz syncline) come from the upper Ludwigsdorf Member and the Lusatiops Member. The taxonomic content of both members is shown in Table 1. The fauna of the upper Ludwigsdorf Member is dominated by hyoliths and disarticulated poriferans, echinoderms, and chancelloriids (Elicki and Schneider, 1992; Elicki, 1994). Further common faunal elements are trilobites, brachiopods, bivalves, archaeogastropods, and monoplacophorans (Table 1). Rather rare are phosphatic microproblematica and cyanobacteria. The fossil assemblage shows distinct changes in its vertical occurrence. Whereas the oldest strata of the upper Ludwigsdorf Member contain only poorly preserved brachiopod and poriferan
TABLE 1. CAMBRIAN FOSSILS OF THE GÖRLITZ SYNCLINE Predominant taxa Miscellaneous taxa Trilobites: Anabarids: Serrodiscus silesius, Lusatiops lusaticus, Lusatiops sp., Ferralsia saxonica, Tiksitheca korobovi, Cambrotubulus Calodiscus cf. lobatus, aff. Calodiscus lobatus, “Holmia” zimmermanni, decurvatus, C. corniformis “Acanthomicmacca” schwarzbachi Hyolithelminths: Molluscs: Torellella mutila, T. lentiformis, Pojetaia runnegari, Fordilla germanica, F. troyensis, Beshtashella tortilis, Pelagiella Hyolithellus cf. micans subangulata, P. lorenzi, P. aff. adunca, Pelagiella sp., Yuwenia juliana, Y.(?) cf. bentleyi, Obtusoconus sp., Planutenia flectata, P. inclinata, Anabarella australis, Bemella aff. jacutica, Bemella sp., Khairkhania evoluta Cyanobacteria: Obruchevella delicata, Endoconchia angusta, Epiphyton sp. Hyolithes: Problematic fossils: Lenalituus pusillus, Conotheca circumflexa, Microcornus elongatus, M. parvulus, Rhombocorniculum cancellatum, Obliquatheca aldanica, Orthotheca sp., Tchuranitheca curvata, Trapezovitus mirus, Aetholicopalla adnata, Coleoloides Hyolithes divaricatus, Hyolithes sp., Burithes sp., Egdetheca cf. aldanica, typicalis, Microcoryne cephalata, sulcavethids indet. Cambroclavus ludwigsdorfensis, Poriferans: Halkieria sp. Dodecaactinella cynodontota, Eiffelia araniformis, unidentified heteractinid remains Chancelloriids: Allonnia tetrathallis, A. tripodophora, Allonnia sp., Archiasterella hirundo, A. pentactina, Chancelloria primaria, Chancelloria sp.
Indeterminate brachiopods and echinoderms Note: For the brachiopods no taxonomic names are presented here because of the rejection of earlier decisions of the extremely poorly preserved and indeterminate material (Freyer 1981a; Ivar Puura, Tartu University, personal commun.). Sources: Based on Freyer (1977), Elicki (1994, 1996, 1998), and Geyer and Elicki (1995).
Paleontological data from the Early Cambrian and paleobiogeographical implications remains, the following levels are of higher diversity. So, the mentioned oldest faunas are followed by an association of brachiopods, poriferans, and hyoliths (with minor content of trilobites and chancelloriids), which, in turn, is overlain by strata yielding only some problematic remains (echinoderms?). Subsequently, a poriferan-dominated fauna developed, with minor trilobite and chancelloriid remains, which changed into the most fossiliferous level bearing a hyolith-poriferan-echinoderm-trilobite–dominated association with chancelloriids, molluscs, and phosphatic small shelly fossils. The youngest carbonate layer just provided a limited poriferan-brachiopod fauna with some hyoliths and chancelloriids. In contrast to the partly very fossiliferous upper Ludwigsdorf Member, the immediately overlying pure shaly Lusatiops Member contains a depleted fauna consisting of some trilobites, scarce brachiopods, and very rare hyoliths. Paleoecologically, the Görlitz fossil assemblages indicate a normal-marine mesotrophic environment dominated by large numbers of semi-infaunal and infaunal suspension feeders and deposit feeders (hyoliths, bivalves, other molluscs, echinoderms, chancelloriids, trilobites, and brachiopods) followed by filter feeders (poriferans) (Elicki, 2003, 2005). The location of this molluscan-dominated habitat is interpreted as relatively proximal. Changes in the physical outer conditions (sea-level changes, siliciclastic influx, salinity, and ?climate) are assumed to be responsible for changes in the diversity and composition of the assemblages (Elicki and Schneider, 1992). Following the paleontological and the sedimentological data, the Görlitz depositional area can be summarized as an openmarine, moderately wave-influenced, well-oxygenated shelf environment during a transgressive stage under mesotrophic conditions. Restricted-marine, higher salinity areas with distinctly depleted fauna were situated locally in proximal position (Elicki and Schneider, 1992; Elicki, 2005). The fauna from the subsurface Early Cambrian carbonates of the Torgau-Doberlug syncline (Zwethau Formation) is quite different from that of the Görlitz syncline. The most conspicuous faunal elements here are archaeocyathans and calcimicrobes (Freyer and Suhr, 1987, 1992; Elicki and Debrenne, 1993; Elicki, 1999b), forming archaeocyathan-calcimicrobial reef mounds and calcimicrobial carpets (Elicki, 1999b; Wotte, 2004). Shelly fossils are represented by some bradoriids and additional phosphatic small shelly fossils, by a few trilobite remains, some chancelloriid and sponge spicules, as well as by undeterminable disarticulated echinoderms (Sdzuy, 1962; Elicki, 1994). Some of them (e.g., poriferans, tintinnoids, bradoriids) are enriched in distinct layers. The widespread occurrence of archaeocyathans, commonly viewed as adapted to low-nutrient conditions (e.g., Brasier, 1990, 1992; Debrenne and Zhuravlev, 1997), supports the assumption of a rather oligotrophic environment. The predominating mode of food production was by photosynthesis (cyanobacteria) and filter-feeding (archaeocyathans). The scarce shelly fauna was represented by suspension- and deposit-feeders (trilobites, chancelloriids), as well as by spicule-bearing filter-feeders (poriferans). Wherever indications of very shallow-water or higher
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salinity conditions occur (lagoons or back shoal areas; see Elicki, 1999a), the abundance of all organisms decreases dramatically. The highest shelly fossil content is in open-marine areas adjacent to archaeocyathan-calcimicrobial reef mounds and to oolitic shoals (Elicki, 1992; Elicki and Debrenne, 1993; Wotte, 2004). The overwhelming majority of organisms lived epibenthicly. A synthesis of the paleontological and sedimentological data points to shallow-ramp depositional environment divided into oolitic shoals, open-marine reef-mounds, and inter-reef areas, and shallow areas of higher salinity; most proximal, intertidal flats were well developed (Elicki, 1999a,b). The fossil content of the Cambrian sediments in the Torgau-Doberlug syncline is shown in Table 2 (for a complete overview, Middle and ?Upper Cambrian fossils are also listed). BIOSTRATIGRAPHY AND PALEOBIOGEOGRAPHY The fossiliferous Early Cambrian sediments of Germany represent different biostratigraphic levels (Fig. 4). For the identification of the biostratigraphic age, scales based on different taxonomic groups were used, so the problem of a high-resolution correlation between these scales must be considered. Thus, from the Early Cambrian assemblage of the Görlitz syncline several trilobites and additional shelly fossils (Table 1) can be used for biostratigraphy. In contrast, from the Early Cambrian of the Torgau-Doberlug syncline only one rather endemic trilobite is reported, but the rich assemblage of archaeocyathans (Table 2) provides a good biostratigraphic datum. For the Ludwigsdorf Member and the overlying Lusatiops Member of the Görlitz syncline, trilobites indicate that the stratigraphic difference between these two members is negligible (Geyer and Elicki, 1995). The trilobite assemblage, consisting mainly of Ferralsia saxonica, Serrodiscus silesius, Calodiscus cf. lobatus, and Lusatiops lusaticus, marks a biostratigraphic position in the higher Banian of the Atlasian series following the new biostratigraphic scheme for western Gondwana proposed by Geyer and Landing (2004). This stratigraphic level corresponds to the middle to upper Marianian stage in Spain (Liñán et al., 1996; Álvaro et al., 1998) and roughly to the upper Atdabanian to early Botoman Bergeroniellus micmacciformis–Erbiella level of the widely used Siberian scale (Brasier, 1989a). The co-occurrence of some small shelly fossils supports this conclusion. Following the shelly-fossil biostratigraphy, the pseudoconodont Rhombocorniculum cancellatum and the gastropod Pelagiella lorenzi, occurring together with the mentioned trilobites, allow correlation with an interval from the Rhombocorniculum cancellatum zone to the Lapworthella cornu zone (sensu Brasier, 1989a) and the Microcornus parvulus zone (sensu Rozanov and Zhuravlev, 1992), respectively. This interval correlates to horizons of Comley (Ac3–Ac4) and Nuneaton (Purley Shale, both England), Newfoundland (lower Brigus Formation), Sweden (Gislöv Formation), China (lower Qiongzhusian), and Siberia (upper Atdabanian to early Botoman) (Brasier, 1989a; Jiang, 1992). Generally the whole assemblage of small shelly fossils
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Middle Cambrian
CU
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TABLE 2. CAMBRIAN FOSSILS OF THE TORGAU-DOBERLUG SYNCLINE Predominant taxa Miscellaneous taxa Trace fossils: Cruziana semiplicata, C. forcata ?Arthropods: Trilobites: Oxyprymna schloppensis Badulesia tenera, Bailiella cf. emarginata, Conocoryphe gemina, C. palpebralis, C. (Parabailiella) sp., Condylopyge rex, C. regia, Clavagnostus? sp., Dorypyge Hyoliths: sp., Ellipsocephalus incultus, Hypagnostus cf. truncatus, Jincella? cf. sulcata, Hyolithes cf. oelandicus, Orthotheca Micmacca anomocaroides, Olenoides sp., Paradoxides brausei, P. saxonicus, aff. affinis, Orthotheca sp. “P.” cf. insularis, P. cf. pinus, P. aff. asturianus, P.? aff. enormis, P. (Acadoparadoxides) sp., Peronopsis sp., Peronopsella inaequalis, Protolenus Undetermined trace fossils, annulatus, P. cf. annulatus, Parasolenopleura lusatica, Solenopleura picardi, echinoderm remains, and Ornamentaspis? frankenwaldensis helcionellids Brachiopods: Acrothele cf. quadrilineata, Acrothele sp., Eoorthis aff. primordialis, Lingulella ferruginea, Lingulella sp.
Early Cambrian
Archaeocyatha: Cordobicyathus germanicus, Nochoroicyathus sp., Urcyathus perejoni, Degeletticyathus? sp., Inessocyathus freyeri, Afiacyathus paracompositus, Erismacoscinus tainius, Erismacoscinus aff. primus, Retecoscinus aff. guadalquivirensis, Coscinocyathus? sp., Neoloculicyathus magnus, Dictyocyathus stipatus, Protopharetra gemmata, Protopharetra dissuta Cyanobacteria: Epiphyton sp., Girvanella sp., Renalcis sp., Kordephyton sp., Proaulopora cf. glabra, Botomaella sp., Subtifloria (= Botominella) sp.
Trilobites: Dolerolichia pretiosa Bradoriids: Hipponicharion elickii, lipabdominids indet. Chancelloriids: Archiasterella pentactina, A. hirundo, Allonnia tripodophora, A. tetrathallis, Chancelloria sp. Hyolithelminths: Torellella lentiformis, T. curva Anabarids: Tiksitheca licis, Cambrotubulus cf. decurvatus Problematic fossils: Tintinnoidella praecursa, Cambroclavus sp. Halkieria sp.
Undetermined trace fossils, and poriferan and echinoderm remains Note: Cu—Upper Cambrian. Sources: Based on Sdzuy (1957a, 1957b, 1962, 1970), Freyer (1981b), Elicki (1994, 1999b), Elicki and Debrenne (1993), and Gozalo and Hinz-Schallreuter (2002).
from the Görlitz syncline is very typical for Gondwanan higher Early Cambrian strata. In contrast to the Görlitz syncline, trilobites are rather rare in Early Cambrian strata of the Torgau-Doberlug syncline. The only known taxon from the area is represented by the neoredlichiid Dolerolichia pretiosa (Sdzuy, 1962), so that the biostratigraphic value reasoned from comparison with taxonomic relatives is limited. Nevertheless, D. pretiosa points to a lower Early Cambrian age (“probably older than the Görlitz trilobite fauna”; Sdzuy, 1962, p. 1096), which can be correlated to the upper part of the lower Ovetian (E. Liñán, 2005, pers. commun.) corresponding to the middle Issendalenian of Geyer and Landing (2004). Much more biostratigraphic data are provided by archaeocyathans. The assemblage of ten regular species (nine genera)
and four irregular species (three genera) from seven drilling cores (Table 2) is best compared with assemblages from Spain and Morocco (Elicki and Debrenne, 1993). Consequently, the Torgau-Doberlug syncline assemblage allows a biostratigraphic classification to the middle Issendalenian of the Atlasian series of western Gondwana (Geyer and Landing, 2004), which is correlative with the lower Ovetian of Iberia and roughly coeval to the lower Atdabanian stage of Siberia. However, a precise age remains a matter of discussion, as the biostratigraphic significance of the West Gondwanan archaeocyathans is not fully understood and systematic assignments depend more or less on individual systematic concepts. The non-trilobitic shelly fauna and the calcimicrobes from the Torgau-Doberlug syncline do not yield further biostratigraphic information.
Paleontological data from the Early Cambrian and paleobiogeographical implications Iberia
Germany GS
TDS
FF
Languedocian
Languedocian
Berglesh.
Caesaraugustan
Caesaraugustan
Lippertsg.
Celtiberian
Middle Cambrian
West-Gondwana
Delitzsch Triebenreuth Tröbitz
Leonian
Wildenstein
Agdzian Galgenberg
Atlasian
Early Cambrian
Bilbilian Banian
Issendalenian
Corduban
Marianian
Ovetian
Tiefenbach
Charlottenh.
Zwethau
Corduban
Figure 4. Stratigraphic scheme of the German Cambrian correlative to the Iberian (Sdzuy et al., 1999) and the western Gondwana scale (proposed by Geyer and Landing, 2004). The Cambrian of Thuringia, found in two drilling cores, is not included because of a lack of biostratigraphic data (available from only a few very poorly preserved microfossils) for a more precise determination within the Cambrian. Berglesh.—Bergleshof Formation; Charlottenh.—Charlottenhof Formation; FF—Frankenwald Forest; GS—Görlitz syncline; Lippertsg.— Lippertsgrün Formation; TDS—Torgau-Doberlug syncline.
Paleobiogeographically, the German Early Cambrian biota show distinct relations to the Mediterranean area. Thus, the genus Lusatiops is known outside the Görlitz syncline only from the Iberian Chains and, amazingly, from Korea (Geyer and Elicki, 1995). The Görlitz genus Ferralsia is reported from the southern Montagne Noire (France) and the Görlitz species F. saxonica from the Sierra Morena (Spain) (Geyer and Elicki, 1995; Álvaro et al., 1998). In the latter region, as in the Görlitz syncline, Serrodiscus silesius co-occurs with Lusatiops and with Ferralsia. The only known Early Cambrian trilobite genus from the Torgau-Doberlug syncline region, Dolerolichia, is not yet reported elsewhere. It belongs to the neoredlichiids, which are widely distributed in the Mediterranean area. The closest morphological relatives of Dolerolichia come from the Spanish-Moroccan region and from southern China (Sdzuy, 1962). The archaeocyathans from the Torgau-Doberlug syncline were formerly used to conclude a paleobiogeographical Siberia connection (Freyer and Suhr, 1987). However, a critical re-investigation of the material has led to the rejection of such an affinity: the archaeocyathans of the Torgau-Doberlug syncline instead
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show distinct relations to the Sierra Morena area (Spain) and Morocco (Elicki and Debrenne, 1993). The non-trilobitic shelly fauna from the Görlitz syncline is very similar to more or less coeval assemblages along the so-called “Paleotethyan belt” (Brasier, 1989b; Bengtson et al., 1990; Brock et al., 2000). Anabarella australis (helcionellid) and Yuwenia (gastropod) are known only from Australia and Germany (Görlitz syncline) (Runnegar in Bengtson et al., 1990; Elicki, 1994, 1996; Brock et al., 2000; Gubanov et al., 2004a,b). Other common molluscs are Pelagiella, Bemella, Obtusoconus, Fordilla, and Pojetaia (Elicki, 1994, 1996). Beshtashella (gastropod) is described from the Görlitz syncline, Spain, Australia, southern China, and Kazakhstan (Elicki, 1994; Gubanov et al., 2004b). Other shelly microfossils also point to a closer connection of areas that were hitherto assumed as positioned at great distances: the pseudoconodont Rhombocorniculum cancellatum from the Görlitz syncline is known from Avalonia, Baltica, Kazakhstan, Mongolia, southern China, and Siberia in roughly coeval strata (e.g., Walliser, 1958; Landing et al., 1980; Brasier, 1989a; Elicki and Schneider, 1992). Additionally, cambroclaves (Görlitz and Torgau-Doberlug synclines), representing very uncommon phosphatic microproblematica, are reported from Sardinia, Australia, Kazakhstan, and China (Mambetov and Repina, 1979; Bengtson et al., 1990; Jiang, 1992; Elicki and Wotte, 2003). In any case, in addition to some of the trilobites (see above), the paleogeographic distribution of the German small shelly fossils seems to support probable Cambrian relationships of the Mediterranean province to eastern Gondwana (Early Cambrian: Pillola, 1991, 1993; Middle Cambrian: Sdzuy, 1972; Late Cambrian: Shergold et al., 2000). Because of the widespread distribution of the Early–Middle Cambrian taxa, Gubanov (2002) and Gubanov et al. (2004b) concluded that the Cambrian paleocontinents must have been situated closer together than was previously assumed. For the Early Cambrian, in any case, the high degree of accordance of the shelly faunas along the Gondwanan margins suggests an intensive and more or less uninterrupted exchange of taxa (1) within the western Perigondwana realm, (2) along the Gondwana margins, and (3) probably to other paleocontinents. DISCUSSION AND CONCLUSIONS Typical for the Cambrian of the Saxo-Thuringian zone of Perigondwana is a sedimentary gap between the late Neoproterozoic (Ediacaran) and the higher Early Cambrian, as well as the widespread absence of Upper Cambrian (Furongian) strata because of nondeposition and/or intensive uplift-related denudation (Linnemann et al., 2000). Generally, not only for the German successions but also for most of central and southern European Mediterranean, Cambrian strata deposited within separated pullapart basins is the widely accepted model (e.g., Buschmann et al., 1995, 2006; Linnemann et al., 2000, Murphy et al., 2004). However, such an assumption of isolated depocenters is not supported by the paleontological data. The demonstrated paleogeographic relationships of the German Early Cambrian benthic fossil
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assemblages cannot be explained if they lived in isolated sedimentary basins. In contrast, the wide geographical distribution, especially of the German trilobite and archaeocyathan taxa, needs not only open-marine conditions but also long-duration migration paths for faunal exchange between the German and the French, Spanish, Portuguese, and Moroccan depocenters, at least (Sdzuy, 1972; Elicki and Debrenne, 1993; Geyer and Elicki, 1995; Álvaro et al., 1998; Liñán et al., 2004). The trilobites and the small shelly fossils of the Görlitz syncline may further suggest long-distance relationships to the other edge of the Paleotethys (see above). Of course, there are doubtless endemic elements in the Cambrian faunas from the Mediterranean (e.g., Sdzuy, 1972; Álvaro et al., 2003). But these do not indicate distinct paleogeographical isolation. Spatial separation of assemblages (and so of genetic information) can also be caused by other, but rather locally acting factors (e.g., temperature, temperature variations, available nutrients, turbidity, water-currents and circulation, environment stability, substrate consistency, direction of coastlines; Dodd and Stanton, 1990). Genetic alterations (mutations) of species in such “separated” communities (e.g., caused by the generally stronger cosmic radiation and the relatively unprotected exposure of organisms within the shallow-water habitats) may result in speciation of new, but now endemic, elements. Such separating outer factors would selectively affect different groups of an assemblage, which can lead to different speciation rates within the “separated” area and can explain the co-occurrence of endemic and non-endemic taxa. There are recent examples for the occurrence of adjoining biotic assemblages containing endemic and non-endemic elements in shortest distance to one another, on the same shelf, without any geographical separation (Sverdrup et al., 1942; Valentine, 1973). In fact, each province or realm, or any biogeographical area, is characterized by a distinct collection of endemic species, and the borders between those regions are usually distinct. A second argument for deposition of the German Cambrian sediments within limited basins should be the seemingly local occurrence of the former in contrast to the succeeding widely distributed Ordovician strata (Linnemann and Buschmann, 1995; Linnemann and Schauer, 1999; Linnemann, 2003). However, some widespread late Middle Cambrian to Early Ordovician phases of uplift and denudation are indicated for different regions of western Perigondwana (e.g., Courtessole, 1973; Bechstädt and Boni, 1994; Linnemann and Buschmann, 1995; Gutiérrez-Marco et al., 2002; Leone et al., 2002). So, not only for Germany but also for the Cambrian succession in Sardinia, for instance, significant erosion of sediments is reported (in Sardinia up to ~2000 m; Bechstädt and Boni, 1994). That level of erosion suggests that the intensities of uplift and denudation have been regionally very different. A further argument against Cambrian deposition in regionally limited basins could be the amazing accordance and continuity of coeval sedimentary facies patterns. Very similar sedimentary successions from Morocco via Portugal, Spain, France, Sardinia, and Germany up to Turkey are highly comparable, which means they are similar over the whole of central Perigondwana (= European
shelf sensu Courjault-Radé et al., 1992) (Dean and Monod, 1970; Bechstädt et al., 1988; Courjault-Radé, 1988; Courjault-Radé et al., 1991; Pillola, 1993; Bechstädt and Boni, 1994; Elicki, 1994; Liñán et al., 1996, 2002; Álvaro et al., 1999; Fernández-Remolar, 1999; Göncüoglu and Kozlu, 2000; Sarmiento et al., 2001). The synthesis of these three points (very active faunal exchange, successively strong but regionally different uplift and denudation, and widely consistent facies realms) supports a modified model of a rather consistent and simply more or less differentiated European shelf of Gondwana, probably containing some areas of higher subsidence (intra-shelf basins). The latter and the different degree of erosion explain the sometimes impressive thicknesses of the preserved sedimentary successions. Thus, the currently interpreted central and southern European terranes represent fragmentary preserved areas of an originally rather consistent central peri-Gondwanan shelf in the Cambrian—a model that may also be extrapolated to the Middle Cambrian (Sdzuy, 1957b, 1962, 1970, 2000; Sdzuy et al., 1999). The content, the coinciding evolution patterns, and the paleogeographic relationships of Perigondwana’s Early to Middle Cambrian faunas demonstrate that a separation and evolution of terranes had not started before the end of the Cambrian or the start of the Early Ordovician. For the Early and Middle Cambrian times at least, a slightly differentiated shelf configuration of central Perigondwana without isolated areas fits best with the paleontological and sedimentological data. ACKNOWLEDGMENTS Many thanks for useful help in the field, for many discussions, and for critical remarks go to Gerd Geyer (Würzburg, Germany), Eladio Liñán (Zaragoza, Spain), Andrey Zhuravlev (Moscow, Russia), and Thomas Wotte and Bernd Buschmann (both Freiberg, Germany). The work was generously supported by the German Research Foundation (research project EL 144/1, 144/12 and Schn 408/1). REFERENCES CITED Álvaro, J.J., Liñán, E., and Vizcaïno, D., 1998, Biostratigraphical significance of the genus Ferralsia (Lower Cambrian, Trilobita): Geobios, v. 31, no. 4, p. 499–504, doi: 10.1016/S0016-6995(98)80121-6. Álvaro, J.J., Vizcaïno, D., and Vennin, E., 1999, Trilobite diversity patterns in the Middle Cambrian of southwestern Europe: A comparative study: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 151, p. 241–254, doi: 10.1016/S0031-0182(99)00033-4. Álvaro, J.J., Elicki, O., Geyer, G., Rushton, A.W.A., and Shergold, J.H., 2003, Palaeogeographical controls on the Cambrian trilobite immigration and evolutionary patterns reported in the western Gondwana margin: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 195, no. 1–2, p. 5–35, doi: 10.1016/S0031-0182(03)00300-6. Bechstädt, T., and Boni, M., 1994, Controls on the evolution of the Cambrian carbonate platform in Bechstädt, T., and Boni, M., eds., Sedimentological, stratigraphical and ore deposits field guide of the autochthonous CambroOrdovician of southwestern Sardinia: Memorie descrittive della carta geologica d’Italia, v. XLVIII: Rome, Servizio Geologico Nazionale, 434 p. Bechstädt, T., Schledding, T., and Selg, M., 1988, Rise and fall of an isolated, unstable carbonate platform: The Cambrian of southwestern Sardinia: Geologische Rundschau, v. 77, no. 2, p. 389–416, doi: 10.1007/BF01832387.
Paleontological data from the Early Cambrian and paleobiogeographical implications Bengtson, S., Conway Morris, S., Cooper, B.J., Jell, P.A., and Runnegar, B.N., 1990, Early Cambrian fossils from South Australia: Brisbane, Association of Australasian Palaeontologists Memoir 9, 364 p. Blumenstengel, H., 1980, Zur Paläontologie und Biostratigraphie der Heinersdorfer Serie (Kambrium) im Thüringer Schiefergebirge: Freiberger Forschungshefte, v. C348, p. 61–68. Brasier, M.D., 1989a, Towards a biostratigraphy of the earliest skeletal biotas, in Cowie, J.W., and Brasier, M.D., eds., The Precambrian-Cambrian boundary: Oxford, Clarendon Press, p. 117–165. Brasier, M.D., 1989b, China and the Palaeotethyan belt (India, Pakistan, Iran, Kazakhstan, and Mongolia), in Cowie, J.W., and Brasier, M.D., eds., The Precambrian-Cambrian boundary: Oxford, Clarendon Press, p. 40–74. Brasier, M.D., 1990, Nutrients in the early Cambrian: Nature, v. 347, no. 6293, p. 521–522, doi: 10.1038/347521b0. Brasier, M.D., 1992, Nutrient-enriched waters in the early skeletal fossil record: Journal of the Geological Society of London, v. 149, p. 621–629. Brause, H., 1969, Das verdeckte Altpaläozoikum der Lausitz und seine regionale Stellung: Abhandlungen der Deutschen Akademie der Wissenschaften, Klasse Bergbau, Hüttenwesen und Montangeologie, v. 1968, no. 1, p. 1–143. Brock, G.A., Engelbretsen, M.J., Jago, J.B., Kruse, P.D., Lauri, J.R., Shergold, J.H., Shi, G.R., and Sorauf, J.E., 2000, Palaeobiogeographic affinities of Australian Cambrian faunas, in Wright, A.J., Young, G.C., Talent, J.A., and Lauri, J.R., eds., Palaeobiogeography of Australian faunas and floras: Brisbane, Association of Australasian Palaeontologists Memoir 23, p. 1–61. Buschmann, B., Linnemann, U., Schneider, J., and Süss, T., 1995, Die cadomische Entwicklung im Untergrund der Torgau-Doberluger Synklinale: Zeitschrift für geologische Wissenschaften, v. 23, no. 5–6, p. 729–749. Buschmann, B., Elicki, O., and Jonas, P., 2006, The Cadomian unconformity in the Saxo-Thuringian Zone, Germany: Palaeogeographic affinities of Ediacaran (terminal Neoproterozoic) and Cambrian strata: Precambrian Research, v. 147, p. 387–403. Courjault-Radé, P., 1988, Analyse sédimentologique de la formation de l’Orbiel (“alternance gréso-calcaires” autc., Cambrian inférieur). Evolution tectono-sédimentaire et climatique (versant sud de la Montagne Noire, Massif central, France): Bulletin de la Société Géologique de France, v. 8, no. 6, p. 1003–1013. Courjault-Radé, P., Gandin, A., Debrenne, F., and Doré, F., 1991, Geodynamic control on Lower Cambrian sedimentation in Normandy, Montagne Noire (France) and Sardinia (Italy), Geologia del Basamento Italiano—Convegno in memoria di Tommaso Cocozza, Sienna, March 21–22, 1991, Abstracts: Siena, Università degli studi, p. 145–147. Courjault-Radé, P., Debrenne, F., and Gandin, A., 1992, Palaeogeographic and geodynamic evolution of the Gondwana continental margins during the Cambrian: Terra Nova, v. 4, p. 657–667. Courtessole, R., 1973, Le Cambrien moyen de la Montagne Noire—Biostratigraphie: Imprim d’Oc, Toulouse, 248 p. Dean, W.T., and Monod, O., 1970, The lower Palaeozoic stratigraphy and faunas of the Taurus Mountains near Beysehir, Turkey—I. Stratigraphy: Bulletin of the British Museum (Natural History): Geology, v. 19, no. 8, p. 413–426. Debrenne, F., and Zhuravlev, A.Yu., 1997, Cambrian food web: A brief review: Geobios, no. 20, p. 181–188. Dodd, J.R., and Stanton, R.J., Jr., 1990, Paleoecology—Concepts and applications: New York, Chichester, Brisbane, Toronto, Singapore, Wiley & Sons, 502 p. Elicki, O., 1992, Faziesanalyse der unterkambrischen Karbonate Deutschlands [Ph.D. thesis]: Freiberg, Freiberg University, 121 p. Elicki, O., 1994, Lower Cambrian carbonates from eastern Germany: Palaeontology, stratigraphy and palaeogeography: Neues Jahrbuch für Geologie und Paläontology, Abhandlungen, v. 191, no. 1, p. 69–93. Elicki, O., 1996, Die Gastropoden und Monoplacophoren der unterkambrischen Görlitz Fauna: Freiberger Forschungshefte, v. C464, p. 145–173. Elicki, O., 1997, Biostratigraphic data of the German Cambrian—Present state of knowledge: Freiberger Forschungshefte, v. C466, p. 155–165. Elicki, O., 1998, First report of Halkieria and enigmatic globular fossils from the central European Marianian (Lower Cambrian, Görlitz syncline, Germany): Revista Española de Paleontología, nùmero extraordinario, Homenaje al Prof. Gonzalo Vidal, p. 51–64. Elicki, O., 1999a, Beitrag zur Lithofazies und zur Lithostratigraphie im Unterkambrium von Doberlug-Torgau: Freiberger Forschungshefte, v. C481, p. 107–119. Elicki, O., 1999b, Palaeoecological significance of calcimicrobial communities during ramp evolution: An example from the Lower Cambrian of Germany: Facies, v. 41, p. 27–40.
151
Elicki, O., 2003, Das Kambrium Sachsens: Veröffentlichungen des Museums für Naturkunde Chemnitz, v. 26, p. 41–62. Elicki, O., 2005, The utility of late Early to Middle Cambrian small shelly fossils from the western Mediterranean: Geosciences Journal, v. 9, p. 161–171. Elicki, O., and Debrenne, F., 1993, The Archaeocyatha of Germany: Freiberger Forschungshefte, v. C450, p. 3–40. Elicki, O., and Schneider, J., 1992, Lower Cambrian (Atdabanian/Botomian) shallow-marine carbonates of the Görlitz syncline (Saxony/Germany): Facies, v. 26, p. 55–66. Elicki, O., and Wotte, Th., 2003, Cambroclaves from the Cambrian of Sardinia (Italy) and Germany: Constraints for the architecture of western Gondwana and the palaeogeographical and paleoecological potential of cambroclaves: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 195, p. 55–71, doi: 10.1016/S0031-0182(03)00302-X. Fernández-Remolar, D.C., 1999, Las calizas fosforíticas del Ovetiense Inferior de la Sierra de Córdoba, España: Boletín de la Real Sociedad Española de Historia Natural, Sección Geológica, v. 95, p. 15–45. Franke, W., 1984, Variszischer Deckenbau im Raume der Münchberger Gneismasse—abgeleitet aus der Fazies, Deformation und Metamorphose im umgebenden Paläozoikum: Geotektonische Forschungen, v. 68, p. 1–253. Freyer, G., 1977, Das Unterkambrium von Görlitz, seine geologische Erforschung und Fossilführung—ein Überblick: Abhandlungen und Berichte des Naturkundemuseums Görlitz, v. 51, no. 7, p. 18–29. Freyer, G., 1981a, Die unterkambrische Brachiopodenfauna des Görlitzer Schiefergebirges: Abhandlungen und Berichte des Naturkundemuseums Görlitz, v. 54, no. 5, p. 1–20. Freyer, G., 1981b, Die Cruzianen von Hainichen bei Borna und das Alter ihrer Fundschichten: Freiberger Forschungshefte, v. C363, p. 51–55. Freyer, G., and Suhr, P., 1987, Über Archaeocyathinen-Funde und den lithologischen Aufbau des Unterkambriums im Gebiet von Torgau: Zeitschrift für geologische Wissenschaften, v. 15, no. 6, p. 665–680. Freyer, G., and Suhr, P., 1992, Über Algengemeinschaften aus der unterkambrischen Zwethauer Folge des Gebietes von Torgau: Freiberger Forschungshefte, v. C445, p. 49–65. Geyer, G., and Elicki, O., 1995, Lower Cambrian trilobites from the Görlitz Synclinorium (Germany)—Review and new results: Paläontologische Zeitschrift, v. 69, no. 1–2, p. 87–119. Geyer, G., and Landing, E., 2004, A unified Lower–Middle Cambrian chronostratigraphy for West Gondwana: Acta Geologica Polonica, v. 54, no. 2, p. 179–218. Göncüoglu, M.C., and Kozlu, H., 2000, Early Paleozoic evolution of NW Gondwanaland: Data from southern Turkey and surrounding regions: Gondwana Research, v. 3, p. 315–323, doi: 10.1016/S1342937X(05)70290-2. Göthel, M., 2001, Das autochthone und allochthone Paläozoikum des “Görlitzer Schiefergebirges” (Mitteleuropäische Varisziden, Deutschland): Zeitschrift für geologische Wissenschaften, v. 29, no. 1–2, p. 55–73. Gozalo, R., and Hinz-Schallreuter, I., 2002, Biostratigraphy and palaeobiogeography of the Cambrian genus Hipponicharion (Ostracoda): Paläontologische Zeitschrift, v. 76, no. 1, p. 65–74. Gubanov, A.P., 2002, Early Cambrian palaeogeography and the probable Iberia-Siberia connection: Tectonophysics, v. 352, p. 153–168, doi: 10.1016/ S0040-1951(02)00194-4. Gubanov, A.P., Skovsted, C.B., and Peel, J.S., 2004a, Anabarella australis (Mollusca, Helcionelloida) from the Lower Cambrian of Greenland: Geobios, v. 37, p. 719–724, doi: 10.1016/j.geobios.2003.05.009. Gubanov, A.P., Fernández-Remolar, D.C., and Peel, J.S., 2004b, Early Cambrian molluscs from Sierra de Córdoba (Spain): Geobios, v. 37, p. 199– 215, doi: 10.1016/j.geobios.2003.04.003. Gutiérrez-Marco, J.C., Robardet, M., Rábano, I., Sarmiento, G.N., San José Lancha, M.Á., Araújo, P.H., and Pieren Pidal, A.P., 2002, Ordovician, in Gibbons, W., and Moreno, T., eds., The geology of Spain: London, Geological Society of London, p. 31–49. Jiang, Z., 1992, The Lower Cambrian fossil record of China, in Lipps, J.H., and Signor, P.W., eds., Origin and early evolution of the metazoan: New York and London, Plenum Press, p. 311–333. Jonas, P., Buschmann, B., and Gaitzsch, B., 2000, Unterkambrischer und unterkarboner Vulkanismus der Torgau-Doberluger Synklinale (NE Saxothuringische Zone): Zeitschrift für geologische Wissenschaften, v. 28, no. 1–2, p. 157–175. Kemnitz, H., Romer, R.L., and Oncken, O., 2002, Gondwana break-up and the northern margin of the Saxothuringian belt (Variscides of central Europe):
152
Elicki
International Journal of Earth Sciences, v. 91, p. 246–259, doi: 10.1007/ s005310100209. Landing, E., Nowlan, G.S., and Fletcher, T.P., 1980, A microfauna associated with Early Cambrian trilobites of the Callavia zone, northern Antigonish Highlands, Nova Scotia: Canadian Journal of Earth Sciences, v. 17, p. 400–418. Leone, F., Ferretti, A., Hammann, W., Loi, A., Pillola, G.L., and Serpagli, E., 2002, A general view on the post-Sardic Ordovician sequence from SW Sardinia: Rendiconti della Società Paleontologica Italiana, v. 1, p. 51–68. Liñán, E., Gámez Vintaned, J.A., and Gozalo, R., 1996, II. Field Conference of the Cambrian Stage Subdivision Working Groups, International Subcommission on Cambrian Stratigraphy, Zaragoza, Spain, September 13–21, Field trip guide and abstracts: Zaragoza, Spain, Universidad de Zaragoza, 129 p. Liñán, E., Gozalo, R., Palacios, T., Gámez Vintaned, J.A., Ugidos, J.M., and Mayoral, E., 2002, Cambrian, in Gibbons, W., and Moreno, T., eds., The geology of Spain: London, Geological Society of London, p. 17–29. Liñán, E., Perejón, A., Gozalo, R., Moreno-Eiris, E., and de Oliveira, J.T., 2004, The Cambrian system of Iberia: Madrid, Publicaciones del Instituto Geológico y Minero de España, Ser. Cuadernos del Museo Geominero, v. 3, 63 p. Linnemann, U., 2003, Sedimentation und geotektonischer Rahmen der Bekkenentwicklung im Saxothuringikum (Neoproterozoikum–Unterkarbon): Geologica Saxonica, v. 48/49, p. 71–110. Linnemann, U., and Buschmann, B., 1995, Die cadomische Diskordanz im Saxothuringikum (oberkambrisch-tremadocische overlap-Sequenzen): Zeitschrift für geologische Wissenschaften, v. 23, no. 5–6, p. 707–727. Linnemann, U., and Schauer, M., 1999, Die Entstehung der Elbezone vor dem Hintergrund der cadomischen und variszischen Geschichte des Saxothuringischen Terranes—Konsequenzen aus einer abgedeckten geologischen Karte: Zeitschrift für geologische Wissenschaften, v. 27, no. 5–6, p. 529–561. Linnemann, U., Gehmlich, M., Tichomirowa, M., Buschmann, B., Nasdala, L., Jonas, P., Lützner, H., and Bombach, K., 2000, From Cadomian subduction to Early Palaeozoic rifting: The evolution of Saxo-Thuringia at the margin of Gondwana in the light of single zircon geochronology and basin development (central European Variscides, Germany), in Franke, W., Haak, V., Oncken, O., and Tanner, D., eds., Orogenic processes: Quantification and modelling in the Variscan belt: London, Geological Society of London, Special Publication 179, p. 131–153. Mambetov, A.M., and Repina, L.N., 1979, The Lower Cambrian of the Talasskiy Alatau and its correlation with the sections of Maly Karatau and the Siberian platform, in Zhuravleva, I.T., and Meshkova, N.P., eds., Biostratigraphy and palaeontology of the Lower Cambrian of Siberia: Moscow, Nauka, p. 98–138 (in Russian). Murphy, J.B., Pisarevsky, S.A., Nance, R.D., and Keppie, J.D., 2004, Neoproterozoic–Early Paleozoic evolution of peri-Gondwanan terranes: Implications for Laurentia-Gondwana connections: International Journal of Earth Sciences, v. 93, p. 659–682, doi: 10.1007/s00531-004-0412-9. Nance, R.D., and Murphy, J.B., 1994, Contrasting basement isotopic signatures and the palinspastic restoration of peripheral orogens: Example from the
Neoproterozoic Avalonian-Cadomian belt: Geology, v. 22, p. 617–620, doi: 10.1130/0091-7613(1994)022<0617:CBISAT>2.3.CO;2. Pillola, G.L., 1991, Trilobites du Cambrien inférieur du SW de la Sardaigne, Italie: Palaeontographica Italica, v. 78, p. 1–174. Pillola, G.L., 1993, The Lower Cambrian trilobite Bigotina and allied genera: Palaeontology, v. 34, no. 4, p. 855–881. Rozanov, A.Yu., and Zhuravlev, A.Yu., 1992, The Lower Cambrian fossil record of the Soviet Union, in Lipps, J.H., and Signor, P.W., eds., Origin and early evolution of the Metazoa: New York, London, Plenum Press, p. 205–309. Sarmiento, G.N., Fernández-Remolar, D., and Göncüoglu, M.C., 2001, Cambrian small shelly fossils from the Cal Tepe Formation, Taurus Mountains, Turkey: Coloquios de Paleontología, v. 52, p. 117–134. Schwarzbach, M., 1934, Das Cambrium der Oberlausitz: Abhandlungen der Naturforschenden Gesellschaft Görlitz, v. 32, no. 2, p. 7–54. Sdzuy, K., 1957a, Revision der mittelkambrischen Trilobiten von Doberlug: Senckenbergiana Lethaea, v. 38, p. 7–28. Sdzuy, K., 1957b, Alter und tiergeographische Stellung des Mittelkambriums von Doberlug: Geologie, v. 6, no. 5, p. 465–475. Sdzuy, K., 1962, Neue kambrische Fossilien aus Bohrungen im Gebiet von Doberlug: Geologie, v. 11, no. 9, p. 1087–1101. Sdzuy, K., 1970, Mittelkambrische Fossilien aus neuen Bohrungen bei Doberlug und bei Delitzsch: Geologie, v. 19, no. 9, p. 1066–1085. Sdzuy, K., 1972, Das Kambrium der acadobaltischen Faunenprovinz—Gegenwärtiger Kenntnisstand und Probleme: Zentralblatt für Geologie und Paläontologie, Teil II, v. 1972, no. 1–2, p. 1–91. Sdzuy, K., 2000, Das Kambrium des Frankenwaldes: Senckenbergiana Lethaea, v. 79, p. 301–327. Sdzuy, K., Liñán, E., and Gozalo, R., 1999, The Leonian stage (early Middle Cambrian): A unit for Cambrian correlation in the Mediterranean subprovince: Geological Magazine, v. 136, p. 39–48, doi: 10.1017/S0016756899002241. Shergold, J.H., Feist, R., and Vizcaino, D., 2000, Early Late Cambrian trilobites of Australo-Sinian aspect from the Montagne Noire, southern France: Palaeontology, v. 43, p. 599–632, doi: 10.1111/1475-4983.00142. Sverdrup, H.V., Johnson, M.W., and Fleming, R.H., 1942, The oceans: Their physics, chemistry, and general biology: Englewood Cliffs, New Jersey, Prentice-Hall, 1087 p. Torsvik, T.H., Smethurst, M.A., Meert, J.G., Van der Voo, R., McKerrow, W.S., Brasier, M.D., Sturt, B.A., and Walderhaug, 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. Valentine, J.W., 1973, Evolutionary paleoecology of the marine biosphere: Englewood Cliffs, New Jersey, Prentice-Hall, 511 p. Walliser, O.H., 1958, Rhombocorniculum comleyense n. gen., n. sp: Paläontologische Zeitschrift, v. 32, p. 176–180. Wotte, Th., 2004, Nachweis und fazielle Position von reef mounds im Unterkambrium von Doberlug-Torgau: Neues Jahrbuch für Geologie und Paläontologie, Abhandlungen, v. 233, no. 3, p. 397–422. MANUSCRIPT ACCEPTED BY THE SOCIETY 3 OCTOBER 2006
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Geological Society of America Special Paper 423 2007
The Variscan orogeny in the Saxo-Thuringian zone—Heterogenous overprint of Cadomian/Paleozoic Peri-Gondwana crust U. Kroner* T. Hahn Institut für Geologie, TU Bergakademie Freiberg, Bernhard-von-Cotta Strasse 2, D-09596 Freiberg, Germany Rolf L. Romer GeoForschungsZentrum Potsdam, Telegrafenberg, D-14473 Potsdam, Germany Ulf Linnemann Staatliche Naturhistorische Sammlungen Dresden, Museum für Mineralogie und Geologie, Königsbrücker Landstrasse 159, D-01109 Dresden, Germany ABSTRACT The Saxo-Thuringian zone of the European Variscides contains the record of the Cadomian and Variscan orogenies and a Paleozoic marine transition stage. The classical view of a relatively simple, double-vergent folded sedimentary basin at the end of the Early Carboniferous is challenged by the widespread occurrence of Late Devonian to Early Carboniferous high-pressure metamorphic units tectonically juxtaposed with low-grade Paleozoic successions. Here we demonstrate that the subdivision of the Saxo-Thuringian zone in three principal units (autochthonous domain, wrench and thrust zone, and allochthonous domain) and their heterogeneous overprint by two regional deformation events during the Variscan orogeny explain the entire geological record. Late Devonian to Early Carboniferous subduction of continental crust inside the allochthonous domain affected a Cadomian basement and sediments deposited on the same continental shelf as the one preserved in the autochthonous domain. Strain partitioning during this regional D1 process led to the formation and evolution of a wrench and thrust zone surrounding the autochthonous domain. The latter was only affected by regional D2 deformation, which was related to regional dextral transpression, rapid exhumation of the subducted rocks of the allochthonous domain, and final filling and subsequent folding of the Saxo-Thuringian flysch basin that covers the autochthonous domain and the wrench and thrust zone. The SaxoThuringian zone is interpreted as a fragment of Peri-Gondwana that never separated from Gondwana to move as an independent terrane and that borders to the Old Red continent, represented by the Rheno-Hercynian zone, along a strike-slip dominated segment of the Rheic suture. The juxtaposition of the Saxo-Thuringian zone with the adjacent areas is discussed as a continuous subduction and/or accretion process representative for the entire Variscan orogen. Keywords: European Variscides, plate tectonics, tectonometamorphic evolution, Saxony, Thuringia *E-mail:
[email protected]. Kroner, U., Hahn, T., Romer, R.L., and Linnemann, U., 2007, The Variscan orogeny in the Saxo-Thuringian zone—Heterogenous overprint of Cadomian/Paleozoic Peri-Gondwana crust, 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. 153–172, doi: 10.1130/2007.2423(06). For permission to copy, contact
[email protected]. ©2007 Geological Society of America. All rights reserved.
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INTRODUCTION The geological record of the Saxo-Thuringian zone is in many respects typical for the continental crust of the European Variscides. The structure of the crust is controlled by at least two orogenies, a Cadomian one and a Variscan one. Sedimentary successions as well as tectonomagmatic and metamorphic events from the latest Precambrian until the Upper Paleozoic witnessed this long-lasting process and led to complex relationships. Previous interpretations were mostly based on the investigation of isolated areas within this region and failed to explain the entire zone in a consistent way. Thus, the increasing number of different, coexisting, and partly contradictory models did not lead to a better understanding of the Saxo-Thuringian zone as a coherent unit. The problematic situation of contrasting and contradictory constraints can be approached in two ways: (1) The area can be subdivided into very small regions of independent pre-Variscan development and called a terrane assemblage (Matte, 1991; Tait et al., 1997; Franke, 2000). (2) A new model explaining the geological data sets from a consistent viewpoint can be established, thereby giving up some of the earlier, model-derived constraints. In this article, we decided to go the second way. We demonstrate that there exists an integrative solution for the Saxo-Thuringian zone and discuss the preferred explanation on a regional scale. The main idea is that all the observed complexities are the result of spatial variations in style and intensity with which the Variscan orogeny affected the common continental crust that had formed in a Peri-Gondwanan setting. Subsequent juxtaposition of the different domains led to the post-Variscan state exposed today. We focus especially on the structural evolution of the different parts of the Saxo-Thuringian zone. In contrast to data from petrology, geochronology, and geochemistry, which led to a reassessment of significant lithological units during the past years, the generally accepted structural geometry has changed little since Kossmat (1927). For example, most of the metamorphic basement rocks earlier believed to be of Precambrian age because of lithostratigraphic considerations (Lorenz and Hoth, 1990) were recognized to be of Lower Carboniferous age (von Quadt, 1993; Kröner et al., 1998; Romer and Rötzler, 2001). Ultra-high pressure minerals in rocks of the Saxo-Thuringian crust (Schmädicke, 1994; Massonne, 1999) require continental subduction, with all its tectonometamorphic consequences. Furthermore, the structural state was explained by long-lasting plane-strain deformation in a general regional northwest-southeast compression, expressing the relative slip vectors of different plates (Franke, 1984, 2000; Oncken, 1997). The observations presented here modify this view and require a rotation of 90° of the proposed direction of initial plate movement. In addition, there are strong arguments against the existence of an independent “Armorican microplate” or the so-called “Armorican terrane assemblage” during the Cambro-Ordovician to Lower Carboniferous. Basin development, isotopic fingerprints of sedimentary rocks, provenance analysis based on detrital zircon grains (e.g., Linnemann et al., 2004), and paleontological
data sets (e.g., Robardet, 2003) require a reinterpretation of the paleogeography. GEOLOGICAL SETTING The Saxo-Thuringian zone of the Variscan orogenic belt in Europe, originally defined by Kossmat (1927), is situated at the northern border of the Bohemian Massif. Traditionally, the faultbound region between the Franconian Line in the southwest and the northeast border of the Lusatian block is considered the main part of the Saxo-Thuringian zone (Fig. 1). The Mid-German crystalline zone, flanking the Saxo-Thuringian zone in the northwest, originally was included in it (Kossmat, 1927), even though it bears the record of an Early Paleozoic subduction-related magmatism (Anthes and Reischmann, 2001), which is absent in other parts of the Saxo-Thuringian zone. The Mid-German crystalline zone separates the Saxo-Thuringian zone from the RhenoHercynian zone, which is located farther to the northwest and comprises predominantly Devonian sediments originally deposited on a passive margin of the Old Red continent (Langenstrassen, 1983). Sedimentary features recording the post-Caledonian sedimentation are well preserved in the Rheno-Hercynian basins, but are missing in units of the Saxo-Thuringian zone (Küstner, 2000). At the southeast margin, a primary fault contact of the Saxo-Thuringian zone against the Teplá Barrandian of the Moldanubian zone is now mostly concealed beneath the Tertiary Eger rift valley. The simplified map (Fig. 1) gives a rough idea of the complex structure of the pre-Permo-Carboniferous basement of the Saxo-Thuringian zone. Crystalline complexes like the Lusatian block, the Erzgebirge, the Granulite Massif, and the Münchberg gneiss complex are juxtaposed on partly unmetamorphosed Paleozoic sediments of the so-called “Saxo-Thuringian basin.” These sediments are believed to build up an uninterrupted sequence of Proterozoic through Late Viséan rocks, whose deposition continued at least up to ca. 330 Ma. This process is supposed to be representative for the whole Saxo-Thuringian zone and for the entire time span (Franke, 1984, 2000). Considering that Variscan subduction and collision-related processes observed in the entire internal orogen, including the Saxo-Thuringian zone, had already started in the Devonian and reached their climax of high-pressure metamorphism at ca. 340 Ma, this model implies that the SaxoThuringian basin is completely sheltered from these orogenic zones until the Late Viséan. The architecture of the middle and lower crust, as derived from geophysical data, seems uniform across the entire region (DEKORP and Orogenic Processes Working Group, 1999). A flat Moho demonstrates the presence of a gravitationally equilibrated crust-mantle boundary at ~30 km depth. Reflector bundles that show a preferred dipping direction toward the southeast were truncated by a laminated lower crust and by the Moho (Behr and Heinrichs, 1987). These features, combined with the vergence of folding inside the Saxo-Thuringian basin, led to the widely used model of a southeast-directed subduction of a Saxo-Thuringian
Figure 1. Simplified tectonostratigraphic map showing the three principal domains of the Saxo-Thuringian zone. Modified after Kroner and Hahn (2004), based on the geological maps from Linnemann and Schauer (1999) and Linnemann and Romer (2002).
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ocean beneath the Moldanubian zone as well as a southeastdirected subduction of a Rheno-Hercynian ocean beneath the Saxo-Thuringian zone. In this model, the final collision is related with northwest-directed nappe transport of Moldanubian allochthonous crystalline complexes over a Saxo-Thuringian foreland. A contemporaneous convergent plate boundary in the northwest is supposed to be responsible for the final collision of the SaxoThuringian zone with the Rheno-Hercynian zone of the Old Red continent. A Rheno-Hercynian retrowedge is postulated for the final southeast-directed stage of the Variscan fold and thrust tectonics inside the Saxo-Thuringian zone (Schäfer et al., 2000). Thus, in this view the Saxo-Thuringian zone represents a small continental terrane acting as a foreland basin of the Moldanubian zone and as a hinterland basin of the Rheno-Hercynian zone (Franke, 1984; Oncken, 1998). Data sets derived from petrological and geochronological investigations of the past decades complicated this regional model. The Saxon granulites and the Erzgebirge, which form the footwall of the postulated autochthonous Saxo-Thuringian basin, do not represent the expected Precambrian basement, but belong to a Variscan high-pressure (HP) to ultrahigh-pressure (UHP) metamorphic unit related to continental subduction and fast exhumation during the Viséan. Pb isotope data demonstrate that the precursors of these HP and UHP complexes represent SaxoThuringian crust rather the Rheno-Hercynian crust, which has a different Pb isotopic signature (Jäckel et al., 1999). This scenario of a subsiding basin that is regionally underlain by contemporaneously subducted and exhumed high-grade metamorphic rocks is called the Saxo-Thuringian paradox (Franke and Stein, 2000). GEOLOGICAL CONSTRAINTS All units of the Saxo-Thuringian zone were affected by the Variscan orogeny. Actually, the entire spectrum of regionally metamorphosed rocks is exposed, ranging from very-low-grade to high-grade rocks, as well as rocks that have experienced ultrahigh temperatures and ultra-high pressures. The lithological, geochemical, and isotopic signatures of the protoliths are relatively uniform. Generally, all units of the Saxo-Thuringian zone preserve features related to the Cadomian orogeny, a Cambro-Ordovician magmatic event, and Paleozoic marine sedimentation. In the autochthonous domain of the Saxo-Thuringian zone most of the pre-Variscan geological record is preserved. Similarly, the metamorphic pile of the Erzgebirge contains the characteristic features of the autochthonous domain. For example, the eastern Erzgebirge contains the complete sequence of protoliths found in the Lusatian block, which was consolidated during the Cadomian orogeny and was affected only slightly by Variscan processes (Tichomirowa et al., 2001). Equivalents of the Paleozoic volcanosedimentary successions of the autochthonous domain of the southeast flank of the Schwarzburg antiform built up different high-pressure tectonometamorphic units in the western Erzgebirge, reaching eclogite-facies conditions (Mingram and Rötzler, 1999). There is no unequivocal evidence for
ophiolithic sequences in the Saxo-Thuringian zone or between the different allochthonous units. The Nd-isotope signature and depleted mantle (TDM) model ages of the Paleozoic marine sediments inside the autochthonous domain indicate that the sediment source did not change from the Late Neoproterozoic until the Lower Carboniferous (pre-flysch deposits); that is, the West African craton was continuously available as a major source supplying detritus to the basin (Linnemann et al., 2004) or sediments derived from it were repeatedly recycled throughout that period. Geochemical signatures reported from the HP metamorphic rocks of the western Erzgebirge, compared to those from the autochthonous domain, reveal that the protoliths principally belong to the same crustal type (Mingram, 1998). Thus, the precursors of the entire Saxo-Thuringian zone seem to belong to the same crust, although there are quantitative differences. For instance, most of the subducted and exhumed HP-UHP orthogneisses in the western Erzgebirge have Ordovician protoliths (Mingram et al., 2004). Such large volumes of Ordovocian magmatic rocks, however, are not observed in the autochthonous domain of the Saxo-Thuringian zone. The pre-Variscan basement rocks of the Saxo-Thuringian zone consist of turbiditic graywacke, volcanoclastics, and plutonic complexes, which are referred to as the “Cadomian basement,” which developed at the Peri-Gondwanan margin of the West African craton between ca. 570 Ma and ca. 540 Ma (Linnemann et al., 2000, 2004; Buschmann et al., 2006). The Cadomian basement formed at an active plate margin at the northern Gondwanan margin, which is termed the “AvalonianCadomian orogenic belt” sensu Nance and Murphy (1994). In general, the basement rocks of this mobile belt are composed of remnants of Neoproterozoic magmatic arc complexes and related marginal basins that developed in a plate-tectonic setting comparable to that of the recent western Pacific region (Buschmann et al., 2001). A slightly different interpretation of the plate-tectonic situation is given by Nance et al. (2002), who explain the geotectonic setting as a Cordilleran-type margin in the Neoproterozoic to the earliest Cambrian. In many regions, the Cadomian magmatic arcs and related intra-arc and back-arc basins were folded and thrusted and became intruded by postkinematic plutons at ca. 540 Ma (Linnemann et al., 2000; Tichomirowa et al., 2001), that is, close to the Precambrian-Cambrian boundary. This stage of deformation and plutonism is usually referred to the so-called “Cadomian orogeny” and is typical for the Avalonian-Cadomian orogenic belt. The timing and setting for the development of the Cadomian basement of the Bohemian Massif and adjacent areas is comparable to that of the type area in the Armorican Massif and other parts of the northern Gondwanan margin (e.g., Iberia, Wales, Atlantic Canada, the Appalachians, the Carolinas, Florida, Mexico; cf. Nance and Murphy, 1996; Nance et al., 2002). The Cadomian unconformity is clearly developed in the Lausitz block, the North Saxon anticline, and the Delitzsch-Torgau-Doberlug syncline, whereas it is not present in the Schwarzburg area (Linnemann
The Variscan orogeny in the Saxo-Thuringian zone and Buschmann, 1995; Linnemann and Romer, 2002). Nance and Murphy (1996) explained the absence of such an unconformity between Neoproterozoic sediments and the Paleozoic overstep sequence at some locations and its presence in others by the relation between the sedimentary basins and the active arc that allowed marginal basins to escape “Cadomian deformation.” The Saxo-Thuringian zone is part of the Peri-Gondwanan province that was affected by Cambro-Ordovician extensional tectonics along the shelf. In the Saxo-Thuringian zone, an inner and outer shelf facies can be distinguished. Lithologies referring to the inner shelf facies occur in the entire Saxo-Thuringian zone, and the type locality is the autochthonous domain (see below). Relicts of the outer shelf facies are preserved in the allochthonous domain. Low-grade metasediments and metavolcanics are overthrust by the crystalline complexes of Münchberg, Wildenfels, and Frankenberg. The characteristic feature of the outer shelf is the extensive Ordovician volcanism. After a period of predominantly clastic sedimentation in the Late Ordovician, a second phase of volcanism started in the Silurian. Increasingly stronger subsidence of the outer shelf facies eventually led in the Silurian and Devonian to deep-water–facies conditions with black shales, pelites, and cherts. Intercalations in late Devonian sediments are dominated by mafic volcanic rocks in the early Frasnian and by felsic volcanic rocks in the Fammenian. In the late Fammenian, sedimentation on the outer shelf changed from pelagic to flysch deposits. Outer shelf facies lithologies are also described from the Bober Katzbach Mountains of the Sudetes (Baranowski et al., 1990) and, farther to the northwest, from the Cambro-Ordovician rift-complex of the Vesser zone (Kemnitz et al., 2002). The HP–UHP units of the Erzgebirge contain metasediments that show lithological similarities with the inner shelf facies and huge volumes of Ordovician metamagmatic rocks that are typical for the outer shelf. Thus, the allochthonous domain contains the whole spectrum of continental crust of the Peri-Gondwanan shelf. The variable overprint of the Saxo-Thuringian crust during the Variscan orogeny, expressed in a great variety of deformational
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structures, reflects different deformation conditions. In a general view, there are two principal regional deformation processes: Regional D1 processes correspond to pervasive deformation during subduction and collision, lasting from 370 to 340 Ma, as indicated by the dating of the metamorphic complexes. D1 processes were always associated with the development of pervasive metamorphic layering. D2 deformation stages refer to ductile deformation related to the final exhumation of HP and UHP metamorphic rocks and their juxtaposition in the upper crust. The model of a widespread Saxo-Thuringian basin with continuous sedimentation lasting at least to 330 Ma precludes D1 processes for this region. Actually, this characteristic is observed only for small parts of the Saxo-Thuringian zone. The occurrence of allochthonous complexes building up large areas of the Saxo-Thuringian basin was first described by Kroner and Hahn (2004). These allochthonous units comprise Ordovician to Lower Carboniferous rocks similar to those described from the autochthonous parts of the basin. Thus, we have to subdivide the Saxo-Thuringian zone into three domains: (1) the autochthonous domain; (2) the wrench and thrust zone; and (3) the allochthonous domain, including all Variscan crystalline complexes. The autochthonous domain and the wrench and thrust zone of the Saxo-Thuringian basin differ by the absence of D1 in the former domain. The characteristic features of the three principal domains of the Saxo-Thuringian zone are summarized in Figure 2. AUTOCHTHONOUS DOMAIN Locally the autochthonous domain shows evidence for a continuous basinal sedimentation, lasting until ca. 330 Ma (Fig. 3) and showing no effects of contemporaneous deformation. The preserved sedimentary and volcanic record in this autochthonous unit covers the time from at least the Ordovician to the Upper Viséan. Early and Middle Cambrian limestones and clastic rocks are only locally preserved, which may be due to restricted primary deposition or erosion during the Late Cambrian.
Figure 2. Features of the three principal domains of the Saxo-Thuringian zone.
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Figure 3. Neoproterozoic to Lower Carboniferous stratigraphy of the unmetamorphosed to very-lowgrade rocks of the Saxo-Thuringian zone. After Baranowski et al. (1990), Gandl (1998), Hahn (1999), Göthel (2001), Linnemann et al. (2004), and Hahn et al. (2005). Numeric time scale from Menning and German Stratigraphic Commission (2002), Ordovician modified according to Gradstein et al. (2004).
The Variscan orogeny in the Saxo-Thuringian zone After a sedimentation gap in the Late Cambrian, continuous marine sedimentation resumed with a transgression in the Early Ordovician (Tremadocian). The Tremadocian sediments of the inner shelf facies consist of up to 3000-m-thick shallow-water siliciclastic deposits, reflecting strong basin subsidence. Felsic and subordinate mafic volcanic and intrusive rocks occur in the basal parts of the sequence. Arenigian to Ashgillian sediments show decreasing sedimentation rates and more distal shelf-facies conditions, dominated by pelites. A gap or strongly condensed sedimentation occurred in the Llanvirnian and Caradocian (horizon of oolithic iron ores). They are overlain by a Late Ordovician glaciomarine tillite reflecting the Sahara glaciation. The Cambro-Ordovician sediments indicate a stage of widespread crustal extension that resulted in a thinned continental crust, already described in adjacent areas of the Bohemian Massif (Kröner et al., 2000; Mazur et al., 2004) as well as in other parts of the European Variscides (Ballevre et al., 2002; Sanchez-Garcia et al., 2003; Laumonier et al., 2004). During the Silurian and Early to Middle Devonian, the inner shelf facies includes undisturbed distal shelf sediments deposited at a low sedimentation rate. Sedimentation was mainly affected by global sea level changes and variations in basin subsidence. Black shales, cherts, deep-water limestones, and pelites were deposited. Thin psammitic intercalations are restricted to the Early Devonian. The Early Frasnian magmatism, which is preserved in the wrench and thrust zone, causes the deposition of turbiditic graywacke with redeposited tuffitic material in the autochthonous domain. This feature demonstrates the close spatial relationship between the autochthonous domain and the wrench and thrust zone. The latter should represent the marginal parts of the basin. The Early Frasnian volcanoclastic rocks are covered by deepwater limestones, with two turbiditic sandstone intercalations in the Late Fammenian. In the Middle Tournaisian, the crenulata event led to the deposition of a black shale, which is overlain by Early Viséan pelites only in the central part of the basin. The onset of distal turbiditic deposits is diachronous, occurring between the Late Tournaisian at the southeast margin and Early Viséan in the central and northwestern parts of the basin. During the Holkerian and the Asbian, distal flysch passed into proximal flysch. A great increase of the sedimentation rate led to the rapid final filling of the basin. After the end of Variscan flysch deposition, the southeastern part of the basin was affected by northwest-directed folding (Franke, 1984). After this local DI deformation the basin was SE–SSE-vergent deformed (DII-deformation). Both deformation processes (DI and DII) belong to the regional D2 deformation event. D1 is absent. WRENCH AND THRUST ZONE This zone shows the same evolution as the autochthonous domain until the Middle Devonian (Fig. 3). The development of the wrench and thrust zone started at the southeast border of the autochthonous domain in the early Frasnian, with subvolcanic
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intrusions and rhyolites. Parts of the subvolcanic granites must have been exhumed rapidly, because granite-conglomerates occur in Frasnian volcanoclastic rocks. After the early felsic phase, voluminous mafic volcanics with pillow lavas were emplaced, followed by and intercalated with tuffites and volcanoclastics. The mafic lavas as well as the associated diabase and picrite dikes show clear within-plate geochemical signatures (Gehmlich et al., 2000) and are mostly related to northeast–southwest-trending faults (Wiefel, 1976; Lange et al., 1999). There is no structural evidence for a rifting event; thus, the rapid exhumation of these felsic intrusive rocks may have been facilitated by the formation of northeast–southwest-oriented strike-slip faults with localized transpression zones. The intense fault-related volcanism resulted in a strong submarine relief that led to facies variations and gaps in the Late Devonian sediments. In the Late Frasnian and Fammenian, sedimentation is dominated by shallow-water to deep-water limestones. Subsequent basin subsidence is indicated by predominantly pelitic deposits in the Late Fammenian and the onset of Variscan distal flysch deposition in the Tournaisian. Distal Variscan flysch was deposited in northeast–southwest-trending basins during the Tournaisian to Early Viséan. In the Arundian (ca. 340 Ma), major parts of the area were involved in Variscan wrench and thrust tectonics (D1 deformation). This deformation was connected with a metamorphic overprint reaching up to greenschist facies. D1 structures comprise thrusts, recumbent folds with NNW to ESE scattering fold axes and subhorizontal foliation planes. Uniformly northeast-trending stretching lineations and pressure shadows around porphyroclasts correspond to the direction of tectonic transport. Kinematic indicators like asymmetric porphyroclasts indicate an initial top-to-thesouthwest transport. The northeast–southwest-trending boundaries between the wrench and thrust zone and the autochthonous domain are considered as D1 strike-slip zones responsible for a decoupling of these two domains and the protection of the autochthonous domain from the regional D1 deformation. Strikeslip faulting may also have been important inside the wrench and thrust zone. This influence is indicated by the juxtaposition of northeast–southwest-striking units comprising similar lithologies with different metamorphic overprint and deformational styles. Because of the subsequent compressional D2 overprint, however, all boundaries now represent reverse faults. The D1-deformation resulted in southwest-directed nappe stacking with shear lenses, duplex structures, and thrusts. An example shows the northeast-southwest profile (Fig. 4) in the Vogtland region (southern wrench and thrust zone). In this area, Paleozoic sequences of the inner shelf facies and distal Lower Carboniferous flysch deposits were detached, partly metamorphosed to greenschist facies, and included in a nappe pile (Hahn, 2003; Hahn and Meinhold, 2005). Especially the less-competent Silurian and Middle Devonian black shales represent preferred detachment zones. The structural character of the imbricated wrench and thrust zone is established by mapping and hundreds of drill holes, cored for uranium exploration
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Figure 4. Northeast-southwest profile showing southwest-directed thrusting in the nappe pile of the southern wrench and thrust zone in the northern Vogtland area. The Lower Carboniferous distal flysch deposits consist of the greenschist facies overprinted Mehltheuer Group and the subgreenschist facies overprinted Elsterberg Group. Both units are roughly coeval, but differ in facies. The Elsterberg Group was overthrust together with Upper Devonian sequences on the Mehltheuer Group and other Paleozoic units.
by the East German SDAG Wismut Company (e.g., Russe, 1995; Lange et al., 1999). Following the end of thrusting and rapid exhumation, wildflysch deposits (Holkerian) and proximal sandstones of the Kahmer Formation (Late Viséan) were unconformably deposited on the nappe pile. The northwest border of the autochthonous domain is built up by similar Lower Paleozoic rock sequences as described above. The Blumenau thrust subdivides the autochthonous domain of the southeast flank of the Schwarzburg antiform from the northern wrench and thrust zone. In contrast to the southeast, rocks reflecting the Late Devonian magmatic event are absent. The observed metamorphic layering as well as northeast-southwest stretched pebbles in the very-low-grade to low-grade metamorphic sequences demonstrate the appearance of the regional D1 deformation inside the northwest flank of the Schwarzburg antiform. In contrast to the crystalline nappes described below, the metamorphic conditions experienced by the allochthonous units of the Saxo-Thuringian basin did not exceed medium pressures and temperatures (Fig. 5). The second deformation event affecting these units (D2) corresponds to the main deformation (DI and DII) recorded in the autochthonous succession of the basin. ALLOCHTHONOUS DOMAIN The allochthonous domain of the Saxo-Thuringian zone includes the crystalline complexes of Münchberg, the Erzgebirge, and the Saxon granulites and low-grade metasedimentary successions. The medium- to high-grade crystalline units of the southeast flank of the Saxo-Thuringian zone, which were partly subjected to extreme P-T conditions (Fig. 5), reflect subduction processes followed by rapid exhumation. From the tectonostratigraphic higher levels to the deeper ones, the following complexes crop out: 1. Eclogite-bearing gneiss complex of Münchberg, gneiss complexes of Wildenfels and Frankenberg; 2. Greenschist and phyllite units partly with lenses of serpentinite; 3. Ordovician low-grade metasediments and metavolcanics, representing the classical example of the outer shelf facies;
4. Low-grade Paleozoic successions similar to the Lower Paleozic sediments and volcanic rocks of the autochthonous domain that form the upper part of the so-called “Fichtelgebirge-Erzgebirge antiform”; 5. The underlying eclogite bearing HP/low-temperature (LT) to UHP/high-temperature (HT) units of the Erzgebirge, representing lithological equivalents to the lower Paleozoic succession of the autochthonous domain (Mingram, 1998); and 6. The lowermost unit of the Erzgebirge, which includes the voluminous medium-pressure (MP) orthogneisses of the eastern Erzgebirge, whose protoliths correspond to the Cadomian basement exposed in the Lusatian block of the autochthonous domain (Tichomirowa et al., 2001). The Saxon granulites are separated from this nappe pile. The granulitic core is mantled by an extremely condensed metamorphic profile. HT shear zones, including migmatitic zones with garnet and cordierite gneisses at the very contact to the HP/UHT granulites, are overlain by shear lenses of serpentinites, strongly deformed gabbros, gneisses, micaschists, and phyllites, and verylow-grade Paleozoic metasediments. The high-grade rocks of this detachment zone experienced a D2 isothermal decompression (Kroner, 1995), the low-grade metasediments were transformed to HT schists at the contact to the hot granulitic core (Reinhardt and Kleemann, 1994). The earliest HP event is recorded in eclogites of the Münchberg Massif, already metamorphosed in the Devonian (Klemd et al., 1991). The metamorphism of the HP complexes of the Erzgebirge took place during the Lower Carboniferous subduction of the continental crust to mantle depth (Schmädicke, 1994; Willner et al., 1997), as documented by findings of microdiamonds (Massonne, 1999). The felsic Saxon granulites, forming a dome structure below the wrench and thrust zone, experienced HP metamorphism at temperatures as high as ~1050 °C (Rötzler and Romer, 2001; Rötzler et al., 2004) at 340 Ma. Tectonic lenses of lithospheric mantle in the Saxon granulites and the UHP units of the Erzgebirge (Mathé, 1969, 1990) provide further evidence for deep subduction of continental lithosphere. The exhumation of the various high-grade complexes to upper crustal levels, which immediately followed the continental subduction, was diachronous within the Saxo-Thuringian zone and
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Figure 5. P-T paths for selected units of the Saxo-Thuringian zone. Note that some high-grade units of the allochthonous domain reached ultrahigh pressure as well as ultra-high temperature conditions. Sources of P-T paths: Brand (1980), Klemd et al. (1991), Reinhardt and Kleemann (1994), Schmädicke (1994), Willner et al. (1997), Kröner and Willner (1998), Massonne (1999), Mingram and Rötzler (1999), Rötzler and Romer (2001), Rötzler et al. (2004). Granite solidus curves according to Stern and Wyllie (1981).
operated in the time span from 400 to 340 Ma (Stosch and Lugmair, 1990; von Quadt, 1993; Kröner and Willner, 1998; Kröner et al., 1998; Romer and Rötzler, 2001, 2003). The uppermost allochthonous units (e.g., Münchberg gneiss complex) arrived in the upper crust at ca. 370 Ma, the high-grade complexes of the Erzgebirge and the Saxon Granulite Massif peaked at 340 Ma and were rapidly exhumed (Romer and Rötzler, 2001). At ca. 330 Ma all high-grade units of the Saxo-Thuringian zone were juxtaposed in the brittle-ductile transition zone of the upper crust (Werner and Lippolt, 2000). This long-lasting tectonometamorphic process (400–340 Ma) produced a very complex structural geometry inside this part of the Saxo-Thuringian zone. The pervasive disturbance of initial lithological contacts is indicated by inverted as well as condensed metamorphic profiles. Recumbent, isoclinal folds, partly only preserved as relics, demonstrate the transpositional character of
the main foliation transecting them. The initial borders of the different inverted metamorphic sequences are oriented parallel to this flat-lying foliation, indicating large-scale thrusting during the process. The evolution of this huge nappe complex and, thus, the metamorphic layering must be diachronous, which is also indicated by the geochronological and stratigraphic constraints for different tectonometamorphic units (Kreuzer et al., 1989; Kröner and Willner, 1998; Werner and Lippolt, 2000). The orientation of mineral-stretching lineations scatters over a range of 90° (Fig. 1). Main azimuths are oriented in northeastsouthwest, east-west, and northwest-southeast directions. This wide variation of the strain increments precludes finite plane strain geometry for the allochthonous domain and demonstrates, together with the coexistence of extensional and compressional structures, the spectacular complexity of this domain. Although the complexity of the structural evolution of the tectonic uplift
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recorded by high-grade metamorphic nappes and mantled gneiss domes is still poorly understood, the exhumation of the complexes probably was attained in two major steps. The initial increments of tectonic exhumation are related to subhorizontal top-to-the-southwest thrusting (Kroner and Hahn, 2004), whereas the final exhumation of the nappes was achieved by subsequent displacements along HT shear zones with top-tothe-west and/or -northwest kinematics. The rapid exhumation of hot continental crust and its juxtaposition against cold supracrustal rocks resulted in HT–very-low-pressure metamorphism (along top-to-the-southeast–directed detachment zones) and the formation of a contact-metamorphic aureole in the roof phyllites enveloping the granulites (Reinhardt and Kleemann, 1994; Kroner, 1995). There is no indication that the exhumation of the Saxon granulites took place beneath an undeformed subsiding sedimentary basin. All rocks above the detachment zone that have been affected by advective heat transfer from the hot granulitic core show evidence for an older regional metamorphism with strong deformation, including recumbent folds and mylonites (Kroner, 1995). The extremely condensed metamorphic profiles belong either to the upper parts of the allochthonous domain or to the phyllites of the wrench and thrust zone. Rapid final exhumation and cooling of the allochthonous domain is contemporaneous with the deposition of the oldest late to post-orogenic sediments. Deposition of early Variscan molasse on top of the uppermost allochthon unit of the Frankenberg Zwischengebirge started at ca. 330 Ma, that is, in the late Asbian (Gehmlich et al., 2000).
SPACE-TIME RELATIONS BETWEEN THE DIFFERENT DOMAINS AND CHARACTERISTICS OF THE SAXO-THURINGIAN BASIN In the three principal domains, the processes of sedimentation, metamorphism, and deformation were not synchronous. The close relationship between these units, however, requires mutual interactions. For instance, the Late Devonian magmatic event in the wrench and thrust zone and the related volcanoclastic sediments in the autochthonous domain prove a close relationship of these two domains during the entire Variscan orogeny. The term Saxo-Thuringian basin is useful to describe sedimentary evolution in the autochthonous domain and the wrench and thrust zone, demonstrating the change from a uniform shelf to a more structured depositional area. The Lower Carboniferous flysch deposits reflect a basin geometry with a northeast–southwesttrending basin axis. The distal-proximal relations of the flysch sediments as well as the direction of the turbiditic currents—with a preferred southwest direction in the axial parts (Gräbe and Wucher, 1967) and northwest-directed sedimentary transport of the wildflysch at the margin to the allochthonous domain—reveal a submarine relief. Thus, we consider the Saxo-Thuringian basin as a Late Devonian to Early Carboniferous syncollisional peripheral foreland basin (sensu Dickinson, 1974) finally bordered by the allochthonous domain in the southeast and the Mid-German crystalline zone in the northwest. During the Viséan, the interplay of deformation and sedimentation between the different domains is obvious (Fig. 6). D1
Figure 6. Mutual interaction of the three principal domains of the Saxo-Thuringian zone in Early Carboniferous time. Rapid overfilling of the Saxo-Thuringian basin and final deformation followed the continental subduction stage of the allochthonous domain and was contemporaneous with the juxtaposition of the three principal domains. Numeric time scale according to Menning and German Stratigraphic Commission (2002).
The Variscan orogeny in the Saxo-Thuringian zone deformation in the wrench and thrust zone is contemporaneous with the onset of proximal flysch deposition at the southeast margin of the autochthonous domain. This proximal flysch comprises redeposited Late Devonian to Tournaisian rocks, shallow-water limestones, and felsic tuffites. Synsedimentary deformation is indicated by widespread unconformable deposition on Upper Devonian to Tournaisian rocks (Gräbe, 1962; Hahn et al., 2005). The increased importance of more proximal flysch deposits in the Holkerian (Middle Viséan) and the shift from proximal flysch to wildflysch deposits in the wrench and thrust zone correlates with the starting exhumation of the Variscan nappe pile (D2 deformation in the wrench and thrust zone). The source area is likely to be farther to the southeast. At this time the uppermost parts of the allochthonous domain must also be closely related to the SaxoThuringian basin (Fig. 7). Variscan flysch deposition ended in the Asbian, when the southeast part of the basin was northwestvergent folded because of the gravitational emplacement of the allochthon crystalline complex of Münchberg (Franke, 1984). Rapid isothermal exhumation of granulites from the subducted parts of the allochthonous domain beneath the wrench and thrust zone, observed at the northwest border of the Saxon granulites, is related to southeast-directed thrusting of the roof phyllites. Advective heat transfer from the deeper parts of the allochthonous domain leads to prograde isobaric heating along the contact to the adjacent wrench and thrust zone. DISCUSSION Causes for Subduction-Exhumation Processes in the HP/ UHP Units of the Allochthonous Domain Subduction processes involving considerable volumes of continental crust have been considered as an important feature of many Phanerozoic collisional belts (e.g., Ernst and Liou, 2000). Single subduction exhumation events are explained by pull of the leading edge of the passive continental margin after the complete consumption of oceanic lithosphere, followed by the buoyant return of the less-dense material (e.g., Chemenda et al., 1997). Multiple subduction exhumation events described for the Scandinavian Caledonides (Brueckner and van Roermund, 2004) require ongoing subduction and exhumation beneath the earliest accreted continental blocks. The latter model is applicable to the data sets of the allochthonous domain and gives an explanation for the long-lasting tectonometamorphic evolution. The uppermost allochthonous unit represents a first Devonian subduction-exhumation event, the Lower Carboniferous HP/UHP units a second one. The final exhumation of subducted units of the Erzgebirge led to the juxtaposition with the older HP units already emplaced in the brittle upper crust. The absence of equivalents of oceanic crust inside the allochthonous pile demonstrates that subduction involved exclusively continental crust during the entire process. One of the big problems in understanding this part of the orogen is the cause for the final juxtaposition of all the different allochthonous units in the Saxo-Thuringian zone. As shown
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Figure 7. Proposed spatial arrangement of the three principal domains of the Saxo-Thuringian zone after the D1 deformation stage at ca. 340 Ma. The uppermost parts of the allochthonous domain are already located in the upper crust, partly reaching the erosion level. The autochthonous domain and the wrench and thrust zone constitute the Saxo-Thuringian flysch basin at this time.
above, there is a connection between the final exhumation of the allochthonous domain and increasing sedimentation in the SaxoThuringian basin. Combined uplift and erosion cannot be the major driving exhumation process for obvious reasons. The intimate tectonic contact between the subducted and exhumed rock complexes with low-grade metasediments in the allochthonous domain—as well as the wrench and thrust zone—requires largescale shear deformation at this stage. Unroofing of the upper crust that was forming metamorphic core complexes during postcollisional regional extensional tectonics discussed for various units of the allochthonous domain (Reinhardt and Kleemann, 1994; Kroner, 1995; Krohe, 1996) explains the existence of localized extensional detachment zones and reduced metamorphic profiles around the rapidly exhumed crystalline complexes. The transition from low-angle extensional ductile shear zones into brittle-ductile and eventually to brittle normal faults (Kroner, 1995) demonstrates the existence of extensional domains during this final juxtaposition process. In this view, the inverted metamorphic profiles across tectonostratigraphic boundaries in the uppermost allochthonous units and the Erzgebirge represent pre-extensional thrusts. Late orogenic extensional collapse (Dewey, 1988) of a complete pre-existing nappe pile does not explain the lenses of coesite- and microdiamond-bearing eclogites and gneisses, respectively, that experienced the metamorphic climax at ca. 340 Ma and that are now exposed as tectonic slices above huge MP orthogneiss complexes. The rapid final exhumation of the deepest parts of the allochthonous domain into mid-crustal levels requires a large amount of thrusting at this stage. Moreover, both in the adjacent wrench and thrust zone and other low-grade units of the more external parts of the mid-European Variscides,
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there are no indications of regional extension in the upper crust between 340 and 330 Ma. Instead this time span is characterized by regional transpressional tectonics along with the development of large-scale ~northwest–southeast-oriented dextral strike-slip faults (Arthaud and Matte, 1977) like the South Armorican shear zone in France (Bosse et al., 2000). The transpressive formation of the fold and thrust belts (e.g., Fielitz, 1992) characteristic for the Variscan orogen outside the internal massifs, starting at this time and persisting at least until the Upper Carboniferous, does not show a spatial relationship to the exhumation processes in the internal zones of the orogen and should represent ongoing convergence of the continental plates during the formation of Pangea. Extrusion tectonics of buoyant crustal slices from a subduction channel (Ernst, 2001; Hynes, 2002, and references therein) could be an appropriate explanation for the tectonometamorphic features of the allochthonous domain (Fig. 8). In such a scenario, where the exhumation path reverses the path of subduction, the subducted units were tectonically juxtaposed with less-deeply buried lithological equivalents. Thrusting at the bottom of the slices is coeval with ductile normal faulting on the top. This is one of the most prominent features of the Erzgebirge as well as the Saxon granulites. The oldest strain increments inside the allochthonous domain, which show an initial northeast-southwest polarity of tectonic transport, therefore should be parallel to the slip vector of the subducted plate during the Variscan convergence. Thus, the long-lasting formation of the allochthonous domain with subsequent subduction-exhumation events should occur in an exclusively convergent plate-tectonic scenario. After 340 Ma, regional metamorphism associated with continental subduction changed to regional HT/LP metamorphism coeval with the final exhumation of the HP/UHP metamorphosed units and was followed by voluminous late orogenic granite magmatism. Thus, the subduction process should have ceased in the Early Carboniferous just before the final exhumation of the deepest parts of the allochthonous domain. In such a scenario, the final west–northwest-directed exhumation during the transpressional stage of the Variscan orogeny represents oblique to lateral extrusion tectonics. The crystalline core of the Saxon granulites with an extensional detachment zone on top should be bound by thrust faults at the bottom (Fig. 9). Deeper parts of the wrench and thrust zone constitute a lateral ramp. Differences in extrusion velocity between individual complexes are the critical parameter for the evolution of ductile to brittle reverse, normal, or perhaps strike-slip faults (Fig. 9). Regional Arguments for a Variscan “Two-Plate” Scenario The occurrence of HP/UHP complexes that formed at different times does not automatically require the existence of different microplates with appended oceanic domains, as discussed in the classical terrane model of the European Variscides. New geochronologic, geochemical, and isotope-geochemical data (e.g., Linnemann et al., 2004) show that the Saxo-Thuringian terrane, a precursor of the pre-Variscan basement of the Saxo-Thuringian
zone, is unlikely to have been an independently drifting microplate, as earlier proposed in several classic interpretations (e.g., Matte, 1991; Franke, 2000). Instead, Saxo-Thuringia was part of Gondwana until the amalgamation of Pangaea during the Carboniferous. Both the Cadomian orogeny and all subsequent preVariscan processes prove a Peri-Gondwanan setting. The striking similarities of paleobiogeographical features of the Paleozoic sediments covering low-grade domains of Iberian, French, and middle European parts of the Variscan orogen (Robardet, 2002, 2003) preclude the closing of oceanic domains south of the Rheic Ocean. Thus, the different domains of the Saxo-Thuringian zone represent part of the Gondwana plate, whereas the RhenoHercynian zone and related areas of Avalonia and the Old Red continent belong to the northern plate. The Upper Ordovician glaciation of Gondwana (Sahara glaciation) is documented in the Saxo-Thuringian zone as a glaciomarine tillite (Ashgillian “Lederschiefer”). Similar tillites are characteristic for all sedimentary domains that remained at the southern margin of the Rheic Ocean (Ghienne, 2003, and references therein). A contemporaneous tillite is not known from Avalonia, which must have drifted away from Gondwana earlier. Thus, the glaciomarine tillite represents an ideal tracer to assign individual crustal fragments to the southern passive margin of the Rheic Ocean, that is, to Gondwana during the Upper Ordovician (Linnemann and Heuse, 2000). However, the Rheno-Hercynian zone spreading across the Anglo-Brabant fold belt over the Rhenish Massif to the MoravoSilesian domain (see Fig. 12 below) is characterized by the Late Caledonian formation of the terrestrial Old Red Sandstone and marine sediments along the adjacent shelf (Langenstrassen, 1983). Detrital mica with a Caledonian age is exclusively found here (Huckriede et al., 2004) and represents the most distinctive feature for continental crust belonging to the northern plate. Because Avalonia and the crust of Peri-Gondwana consist of a Cadomian basement this feature is not appropriate for the delimitation of both plates. Notwithstanding, the former proximitiy of Avalonia to the Amazonia craton during the Cadomian orogeny is documented in Grenville ages of inherited zircon (e.g., Murphy et al., 2004, and references therein) and contrasts with the west African provenance of the Peri-Gondwana crust. The rare occurrence of those zircons in the high-pressure units of the Erzgebirge (Mingram et al., 2004) has been used to argue that parts of the Saxo-Thuringian and Rheno-Hercynian zones have the same evolution and belong to Avalonia. This conclusion contradicts the lithological and geochemical similarities of the allochthonous and the autochthonous domains, discussed above. Furthermore, trace-element signatures of Carboniferous and Permian granitoids from the Rheno-Hercynian and SaxoThuringian zones (Förster and Tischendorf, 1996) and lead isotope data from ore deposits that are spatially and genetically closely related with these granitoids (Bielicki and Tischendorf, 1991) differ systematically. Because the geochemical and Pb isotopic signature of these post-Variscan granitoids and ore deposits is strongly influenced by assimilation of crustal rocks (Förster
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Figure 8. Orogenic wedge model with a subduction channel beneath the prowedge during the Early Carboniferous stages of the D1 evolution until 340 Ma. The orogenic wedge is built up of the Early Variscan subducted and exhumed metamorphic units of the allochthonous domain and the accreted Cadomian block of the Teplá Barrandian. Subduction-exhumation paths of the HP/UHP units of the Erzgebirge and the phase transitions along the channel are derived from the P-T paths of Figure 5. Eclogitic lower crust and lithospheric mantle represent the footwall and the hanging wall, respectively, of the channel. The maximum subduction depth of upper and middle crustal units is limited by melting of dry granitic crust.
Figure 9. Final exhumation of the deeper parts of the allochthonous domain is related to the juxtaposition of the Saxon granulites beneath the wrench and thrust zone and dextral strike-slip shearing between the Erzgebirge and the autochthonous domain of the Lausitz block. Kinematic indicators demonstrate exhumation direction toward the west and northwest. Lateral extrusion tectonics along a lateral ramp (l.r.) in the deeper parts of the wrench and thrust zone is considered for this process. The relative tectonic movement caused by different extrusion velocities between adjacent areas contemporaneously produces extensional, compressional, and strike-slip structures during the regional D2 deformation (<340 Ma). Isothermal exhumation along HT shear zones leads to isobaric heating at the overlying tectonic contact of the wrench and thrust zone caused by advective heat transfer.
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Kroner et al. and Tischendorf, 1996; Romer et al., 2001), the geochemical and isotopic contrast between post-Variscan granites and mineralizations in the Saxo-Thuringian and Rheno-Hercynian zones precludes a common source for the Ordovician to Carboniferous sediments in the two domains. Rather than interpreting rare Grenville ages in pre-Variscan magmatic rocks of the allochthonous domain of the Saxo-Thuringian zone as artifacts, we consider them as an indication for the initial proximity of the leading edge of the thinned Gondwana shelf to Avalonia and as an indication that the subducted continental crust, now forming metamorphic complexes of the Erzgebirge, belongs partly to the outer shelf facies. We propose a subduction-collision model with a subducting lower Gondwana plate, starting with the crust of the Rheic Ocean beneath the upper plate of the Old Red continent. In such a twoplate scenario, the Rheno-Hercynian zone and related areas of the Old Red continent represent the hinterland of the overriding plate from the beginning of subduction in the Silurian. Closure of the Rheic Ocean
Figure 10. Schematic map of the present-day tectonics along the active plate margin of the Sumatra region.
Figure 11. Principal plate-tectonic scenario for the Variscan orogeny in Europe. Clockwise rotation of Gondwana caused by the opening of the Paleotethys results in the closure of the Rheic Ocean and oblique convergence between Peri-Gondwana and the Old Red continent between the Silurian and the Carboniferous. The southern Gondwana plate is the subducted plate, the northern plate is the overriding one.
The Early Variscan convergent tectonics is connected with the subduction of the Rheic Ocean. Because oblique convergence is globally much more common than frontal subduction and collision, we believe that the active plate margin between the eastern Himalayan syntax and the Sumatra region (Fig. 10) could serve as a present-day analogue for the Early Variscan situation (Fig. 11). Oblique convergence between the lower (Indo-Australian) and the upper (Eurasian) plates led to extension and strike-slip tectonics in the upper plate during coeval subduction normal to the plate boundary. This slip partitioning, demonstrated for the earthquake slip vectors along the Sumatra region (McCaffrey, 1996), is responsible for margin parallel strike-slip displacement of the trench and the forearc regions (Malod and Kemal, 1996). Whereas the seafloor spreading in the back-arc basin of the Andaman Sea is controlled by transtensional tectonics along trench parallel strike-slip faults, the initiation of the opening of the Andaman Sea may be either the result of extrusion tectonics during the IndiaAsia collision (Kamesh Raju et al., 2004) or the consequence of the nonconformity between plate shape and subduction margin geometry (Khan and Chakraborty, 2005). Transpressional tectonics before and during the India-Eurasia collision discussed for the Late Cretaceous–Paleogene transpressional belt in the northern region of Sundaland (Morley, 2004), furthermore, demonstrates the synchronism of compressional, extensional, and strike-slip tectonics in such a scenario. Ongoing convergence led to the juxtaposition of the peripheral foreland basins of India with basins of the overriding plate along the active boundary. If we consider an initial northeast-directed plate movement of the Gondwana plate relative to the Old Red continent during the Lower Devonian, as derived from the strain increments of the D1 deformation in the Saxo-Thuringian zone, the following features of the proposed overriding plate can be fully explained by oblique convergence (Fig. 12A). Sinistral transtension in the Old
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Figure 12. Oblique subduction and collision processes along the European segment of the Rheic suture (schematic map). (A) Diachronous accretion of dispersed Cadomian blocks during ongoing oblique subduction of the leading edge of the Gondwana plate is related to spatially restricted compressional, extensional, and strike-slip tectonics. This model explains the Acadian deformation of the Anglo-Brabant fold belt, the extension in the Rheno-Hercynian basin, and the coeval subduction-exhumation event of the uppermost allochthonous domain during the Devonian with the relative movement of two lithospheric plates. LB—Lausitz block; TB—Teplá Barrandian block; 1a—uppermost allochthonous units; 1b—future lowermost units of the allochthonous domain; 2a/2b—southern and northern wrench and thrust zones, respectively; 3—autochthonous domain. (B) Collision of the Teplá Barrandian block with the already-exhumed uppermost allochthonous units and eastern parts of the Rheno-Hercynian basin is related to the starting subduction of continental crust that forms the lowermost units of the allochthonous domain. (C) Ongoing convergence eventually led to the present spatial arrangement of the different Cadomian blocks along the Rheic suture. Also, sedimentary basins belonging to the two different plates are juxtaposed along strike-slip–dominated parts of the Rheic suture. The Bohemian Massif and the MidGerman crystalline zone constitute an orogenic wedge and a sinistral transpressional belt, respectively, and border the Saxo-Thuringian flysch basin. SW—Schwarzwald; MGCZ—Mid-German crystalline zone. (D) The final stage of the Variscan orogeny is characterized by orogen-wide transpressional tectonics, regional HT/LP metamorphism, late orogenic granite intrusions, and the formation of the fold and thrust belts of the external parts of the orogen. Furthermore, the related final extrusion of deeper parts of subducted units belonging to Peri-Gondwana and their juxtaposition in the lithologically equivalent upper crust is an important feature of the Variscan orogen.
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Red Sandstone basins predates Acadian deformation inside the Anglo-Brabant fold belt (Soper and Woodcock, 2003). This compressional event at ca. 400 Ma (e.g., Mansy et al., 1999; Sherlock et al., 2003) is coeval with ongoing subsidence and crustal extension inside the adjacent Rheno-Hercynian basin (Grösser and Dörr, 1986). Strain partitioning inside the upper plate during sinistral oblique convergence could be an appropriate explanation for this complex situation. Thus, the Acadian deformation could be an Early Variscan expression of the first collisional contacts between the Gondwana plate and the Old Red continent in this area, as discussed by Soper and Woodcock (2003), rather than the commonly proposed Late Caledonian event. Moreover there is no need for a Rheno-Hercynian Ocean. We interpret this zone as a localized transtensional domain in a hinterland position that is possibly connected with the evolution of small volumes of oceanic crust inside a short-lived back-arc basin (Leeder, 1982). Contrasting Material Paths of Peri-Gondwana Crust during Variscan Convergence High-pressure metamorphism in allochthonous units surrounding the low-grade crustal blocks of Peri-Gondwana demonstrate a large amount of crustal shortening during continuous convergence from the Devonian to the Early Carboniferous. Continental crust belonging to the Lower plate in part was subducted to mantle depth and emplaced in huge nappe piles, tectonically juxtaposed with the same rocks as those of the adjacent autochthonous domain, which remained in the upper crust during the whole orogeny. This dualism in material path could take its origin in differences in crustal thickness and/or composition of the preVariscan Peri-Gondwanan crust. In the Cadomian massifs of the Variscan orogen—as, for instance, in the region of the British Channel islands (D’Lemos et al., 2001), the north and central Armorican zone (Nagy et al., 2002), the Teplá Barrandian zone of the Bohemian Massif (Dörr et al., 2002), and the autochthonous domain of the Saxo-Thuringian zone (Linnemann et al., 2000)—voluminous late Cadomian magmatic complexes dominate the crystalline basement. In comparison with these massifs, the metamorphic domains of the Variscides additionally are characterized by large volumes of post-Cadomian Cambro-Ordovician magmatic rocks (e.g., Pin et al., 1992; Turniak et al., 2000; Ballèvre et al., 2002; this study). In some regions, the Cadomian basement is completely absent (Laumonier et al., 2004). The precollisional Peri-Gondwanan crust represents a widespread continental extensional province rather than a simple shelf of a rifted passive continental margin. Large-scale crustal thinning possibly led to the Cambro-Ordovician evolution of dispersed arranged continental blocks representing “normal” inner shelf and very thin outer shelf facies crust with large numbers of Paleozoic magmatic rocks. The inner shelf crust forms relatively rigid blocks that accreted together with the early Variscan exhumed HP units at the active margin, when convergence still was ongoing. In contrast, the outer shelf crust seems to have
been subductable. This subduction-accretion process could lead to spatially restricted orogenic wedge geometries (terminology after Johnson and Beaumont, 1995), with the early accreted crustal block in the upper plate position. Ongoing subduction of the thinned continental crust occurred beneath the pro-wedge, whereas the region near the former plate boundary is situated at the retro-wedge. The primary hinterland of the overriding plate now represents a retro-foreland basin (Fig. 12B,C). Crustal shortening in both plates is connected with largescale strike-slip tectonics and localized transpressional and transtensional domains along the lateral limitations of the particular orogenic wedges, which avoids strain incompatibilities. Continuous sedimentation with subsequent deformation occurs in the proand retro-foreland basins. This scenario could be an explanation for the uneven configuration of the European Variscides and the reason for the asynchronous occurrence of similar processes in the different regions of the entire orogen. The Bohemian Massif combined with the adjacent Saxo-Thuringian zone (Fig. 13) is an excellent example demonstrating a more or less complete sequence of the Variscan processes discussed above and represents, from our point of view, a type locality that explains this spectacular orogen in Europe. CONCLUSIONS The proposed evolutionary model implies the following constraints: • Peri-Gondwana represents the northern Lower Paleozoic continental shelf of the Gondwana plate with ongoing marine sedimentation during the Lower Paleozic. It consists of Cadomian crystalline blocks with overlying inner shelf facies sediments surrounded by thinned Cadomian basement and large volumes of magmatites caused by a large-scale Cambro-Ordovician continental extension event related to the opening of the Rheic Ocean and the separation of Avalonia. This post-Cadomian thinned continental crust, with its outer shelf facies Paleozoic cover, is considered to be subductable. • Oblique northeast-directed convergence between the southern Gondwana plate and the northern plate of the Old Red continent started in the Silurian and is connected with the subduction of the Rheic Ocean beneath the northern plate. • Diachronous collision of the dispersed Cadomian blocks is coeval with subduction of the thinned continental crust of the Gondwana plate and persisted from the Early Devonian until the Early Carboniferous (i.e., a time span of ~60 m.y.). • Early accreted continental blocks of the Gondwana plate together with Early Variscan exhumed HP metamorphic units form spatially restricted, fault-bounded orogenic wedges, with ongoing subduction of the Gondwana plate beneath the pro-wedge and coeval thrusting in the retrowedge area of the northern plate.
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Figure 13. Schematic northwest-southeast profiles of the Saxo-Thuringian zone across the northeast–southwest-oriented fault zones that are parallel to the strike-slip–dominated part of the Rheic suture zone. The SaxoThuringian basin formed during regional D1 deformation of the Variscan orogen between the Mid-German crystalline zone and the developing allochthonous domain, which belongs to the Bohemian Massif. Northwest-directed mass transfer from the overthickened allochthonous domain during D2 led to the final crustal architecture of the Saxo-Thuringian zone. Widespread southeast-vergent folding is due to backthrusting during the final transpressional stage of the Variscan orogeny. Symbols of the geological units are the same as in Figure 1.
• Northeast–southwest-directed transform zones with localized transpressional and transtensional domains sustain strain compatibility. • During ongoing convergence, syncollisional foreland basin geometries developed on both plates. On the Gondwana plate they represent pro-foreland basins. In contrast, the basins of the northern plate have a retro-foreland character. Ongoing convergence led to the juxtaposition of these two basin types along strike-slip–dominated segments of the Rheic suture zone. • The final exhumation of the subducted continental crust took place during regional Late Variscan transpressional tectonics. Associated advective heat transfer to upper crustal levels from exhumed rocks is the likely reason for
regional HT/LP metamorphism. During this stage, which is additionally characterized by large-scale northwest– southeast-directed dextral strike-slip zones and voluminous granite magmatism, extrusion tectonics along the former subduction channel led to the juxtaposition of units with quite different evolution histories. ACKNOWLEDGMENTS Funding by the German Science Foundation (Kr 1566/1-1) is gratefully acknowledged. Careful reviews and helpful comments by C. Quesada and an anonymous reviewer are greatly appreciated. L. Ratschbacher is thanked for the fruitful discussion about the present-day tectonics in Southeast Asia.
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REFERENCES CITED Anthes, G., and Reischmann, T., 2001, Timing of granitoid magmatism in the eastern mid-German crystalline rise: Journal of Geodynamics, v. 31, no. 2, p. 119–143, doi: 10.1016/S0264-3707(00)00024-7. Arthaud, F., and Matte, P., 1977, Late Paleozoic strike-slip faulting in southern Europe and northern Africa; result of a right-lateral shear zone between the Appalachians and the Urals: Geological Society of America Bulletin, v. 88, no. 9, p. 1305–1320, doi: 10.1130/0016-7606(1977)88<1305: LPSFIS>2.0.CO;2. Ballèvre, M., Capdevila, R., Guerrot, C., and Peucat, J.-J., 2002, Discovery of an alkaline orthogneiss in the eclogite-bearing Cellier unit (Champtoceaux Complex, Armorican Massif): A new witness of the Ordovician rifting: Comptes Rendus Geosciences, v. 334, p. 303–311, doi: 10.1016/ S1631-0713(02)01760-1. Baranowski, Z., Haydukiewicz, A., Kryza, R., Lorenc, S., Muszynski, A., Solecki, A., and Urbanek, Z., 1990, Outline of the geology of the Gory Kaczawskie (Sudetes, Poland): Neues Jahrbuch für Geologie und Paläontologie Abhandlungen, v. 179, p. 223–257. Behr, H.J., and Heinrichs, T., 1987, Geological interpretation of DEKORP 2–5: A deep seismic reflection profile across the Saxothuringian and possible implications for the late Variscan structural evolution of central Europe: Tectonophysics, v. 142, p. 173–202, doi: 10.1016/00401951(87)90122-3. Bielicki, K.H., and Tischendorf, G., 1991, Lead isotope and Pb-Pb model age determination of ores from central Europe and their metallogenetic interpretation: Contributions to Mineralogy and Petrology, v. 106, p. 440–461, doi: 10.1007/BF00321987. Bosse, V., Feraud, G., Ruffet, G., Ballèvre, M., Peucat, J.-J., and De Jong, K., 2000, Late Devonian subduction and early-orogenic exhumation of eclogite-facies rocks from the Champtoceaux Complex (Variscan belt, France): Geological Journal, v. 35, p. 297–325, doi: 10.1002/gj.864. Brand, R., 1980, Die niedriggradige Metamorphose einer Diabas-Assoziation im Gebiet von Berg/Frankenwald: Neues Jahrbuch für Mineralogie Abhandlungen, v. 139, p. 82–101. Brueckner, H.K., and van Roermund, H.L.M., 2004, Dunk tectonics; a multiple subduction/eduction model for the evolution of the Scandinavian Caledonides: Tectonics, v. 23, no. 2, p. 20. Buschmann, B., Nasdala, L., Jonas, P., Linnemann, U., and Gehmlich, M., 2001, SHRIMP U-Pb dating of tuff-derived and detrital zircons from Cadomian marginal basin fragments (Neoproterozoic) in the northeastern Saxothuringian zone (Germany): Neues Jahrbuch für Geologie und Paläeontologie, Monatshefte, v. 2001, p. 321–342. Buschmann, B., Elicki, O., and Jonas, P., 2006, The Cadomian unconformity in the Saxo-Thuringian zone, Germany: Palaeogeographic affinities of Ediacaran (terminal Neoproterozoic) and Cambrian strata: Precambrian Research, v. 147, no. 3/4, p. 387–403. Chemenda, A., Matte, P., and Sokolov, V., 1997, A model of Palaeozoic obduction and exhumation of high-pressure/low-temperature rocks in the southern Urals: Tectonophysics, v. 276, no. 1–4, p. 217–227, doi: 10.1016/S00401951(97)00057-7. DEKORP and Orogenic Processes Working Groups, 1999, Structure of the Saxonian Granulites: Geological and geophysical constraints on the exhumation of high-pressure/high-temperature rocks in the mid-European Variscan belt: Tectonics, v. 18, p. 756–773, doi: 10.1029/1999TC900030. Dewey, J.F., 1988, Extensional collapse of orogens: Tectonics, v. 7, no. 6, p. 1123–1139. Dickinson, W.R., 1974, Plate tectonics and sedimentation, in Dickinson, W.R., ed., Tectonics and sedimentation: Tulsa, Oklahoma, Society of Economic Paleontologists and Mineralogists Special Publication 22, p. 1–27. D’Lemos, R.S., Miller, B.V., and Samson, S.D., 2001, Precise U-Pb zircon ages from Alderney, Channel Islands: Growing evidence for discrete Neoproterozoic magmatic episodes in northern Cadomia: Geological Magazine, v. 138, no. 6, p. 719–726. Dörr, W., Zulauf, G., Fiala, J., Franke, W., and Vejnar, Z., 2002, Neoproterozoic to Early Cambrian history of an active plate margin in the TepláBarrandian unit—A correlation of U-Pb isotopic-dilution-TIMS ages (Bohemia, Czech Republic): Tectonophysics, v. 352, no. 1–2, p. 65–85, doi: 10.1016/S0040-1951(02)00189-0. Ernst, W.G., 2001, Subduction, ultrahigh-pressure metamorphism, and regurgitation of buoyant crustal slices—Implications for arcs and continental
growth: Physics of the Earth and Planetary Interiors, v. 127, no. 1–4, p. 253–275, doi: 10.1016/S0031-9201(01)00231-X. Ernst, W.G., and Liou, J.G., 2000, Overview of UHP metamorphism and tectonics in well-studied collisional orogens: International Book Series, v. 4, p. 3–19. Fielitz, W., 1992, Variscan transpressive inversion in the northwestern central Rhenohercynian belt of western Germany: Journal of Structural Geology, v. 14, no. 5, p. 547–563, doi: 10.1016/0191-8141(92)90156-Q. Förster, H.-J., and Tischendorf, G., 1996, Compositional heterogeneity of silicic magmatic rocks from the German Variscides: Zeitschrift für geologische Wissenschaften, v. 24, p. 467–482. Franke, W., 1984, Variszischer Deckenbau im Raum der Münchberger Gneismasse—abgeleitet aus der Fazies, Deformation und Metamorphose im umgebenden Paläozoikum: Geotektonische Forschungen, v. 68, p. 1–253. Franke, W., 2000, The mid-European segment of the Variscides: Tectonometamorphic units, terrane boundaries and plate tectonic evolution, in Franke, W., Haak, V., Oncken, O., and Tanner, D., eds., Orogenic processes— Quantification and modelling in the Variscan belt of central Europe: London, Geological Society of London Special Publication 179, p. 35–61. Franke, W., and Stein, E., 2000, Exhumation of high-grade rocks in the SaxoThuringian belt: Geological constraints and geodynamic concepts, in Franke, W., Haak, V., Oncken, O., and Tanner, D., eds., Orogenic processes—Quantification and modelling in the Variscan belt of central Europe: London, Geological Society of London Special Publication 179, p. 337–354. Gandl, J., 1998, Neue Daten zum jüngeren Paläozoikum NE-Bayerns und angrenzender Gebiete—Faziesentwicklung und geotektonische Konsequenzen: Geologica Bavarica, v. 103, p. 19–273. Gehmlich, M., Linnemann, U., Tichomirowa, M., Gaitzsch, B., Kroner, U., and Bombach, K., 2000, Geochronologie oberdevonischer bis unterkarbonischer Magmatite der Thüringischen und Bayerischen Faziesreihe sowie variszischer Deckenkomplexe und der Frühmolasse von Borna-Hainichen (Saxothuringisches Terrane): Zeitschrift der deutschen geologischen Gesellschaft, v. 151, p. 337–363. Ghienne, J.-F., 2003, Late Ordovician sedimentary environments, glacial cycles, and post-glacial transgression in the Taoudeni basin, west Africa: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 189, no. 3–4, p. 117–145, doi: 10.1016/S0031-0182(02)00635-1. Göthel, M., 2001, Das autochthone und allochthone Paläozoikum des Görlitzer Schiefergebirges (Mitteleuropäische Varisziden, Deutschland): Zeitschrift für geologische Wissenschaften, v. 29, p. 55–73. Gräbe, R., 1962, Beziehungen zwischen der tektonischen und faziellen Entwicklung des Oberdevons und Unterkarbons sowie zur Genese der Eisenerze vom Lahn-Dill-Typus am NW-Rand des Bergaer Sattels: Freiberger Forschungsheft, v. C, no. 140, p. 1–83. Gräbe, R., and Wucher, K., 1967, Schüttungs- und Strömungsrichtungen im Kulm des SE-Teils der Ziegenrücker Mulde (Ostthüringisches Schiefergebirge): Geologie, v. 16, p. 991–1006. Gradstein, F.M., Ogg, J.G., and Smith, A.G., 2004, A geologic time scale 2004: Cambridge, Cambridge University Press, 589 p. Grösser, J., and Dörr, W., 1986, MOR-Basalte im östlichen Rheinischen Schiefergebirge: Neues Jahrbuch für Geologie und Paläontologie, Monatshefte, v. 12, p. 705–722. Hahn, T., 1999, Die unterkarbone Sedimentation im östlichen Variszikum— Zusammenfassende Darstellung publizierter Daten [M.S. thesis]: Freiberg, Technische Universität Bergakademie Freiberg, 56 p. Hahn, T., 2003, Geologische Karte von Thüringen, Blatt 5339 Greiz: Jena, Thüringer Landesanstalt für Umwelt und Geologie, scale 1:25,000. Hahn, T., and Meinhold, G., 2005, Revision der Unterkarbonstratigraphie der Mehltheuerer Mulde (Vogtland, Saxothuringikum): Geowissenschaftliche Mitteilungen von Thüringen, v. 12, p. 51–70. Hahn, T., Wucher, K., Heuse, T., and Melzer, P., 2005, Neudefinition lithostratigraphischer Einheiten im Unterkarbon (Kulm) des Thüringisch-Fränkisch-Vogtländischen Schiefergebirges: Geowissenschaftliche Mitteilungen von Thüringen, v. 12, p. 19–49. Huckriede, H., Wemmer, K., and Ahrendt, H., 2004, Palaeogeography and tectonic structure of allochthonous units in the German part of the RhenoHercynian belt (central European Variscides): International Journal of Earth Sciences, v. 93, no. 3, p. 414–431, doi: 10.1007/s00531-004-0397-4. Hynes, A., 2002, Encouraging the extrusion of deep-crustal rocks in collisional zones: Mineralogical Magazine, v. 66, no. 1, p. 5–24, doi: 10.1180/0026 461026610013.
The Variscan orogeny in the Saxo-Thuringian zone Jäckel, P., Anthes, G., Molzahn, M., Oncken, O., and Reischmann, T., 1999, Pb isotopic systematics of the central Variscan belt, in Dietrich, P.G., Franke, W., Merkel, B.J., and Herzig, P., eds., Terra Nostra (Bonn), v. 99-1, p. 111–112. Johnson, D.D., and Beaumont, C., 1995, Preliminary results from a planform model of orogen evolution, surface processes and the development of clastic foreland basin stratigraphy, in Dorobek, S.L., and Ross, G.M., eds., Evolution of foreland basins: Tulsa, Oklahoma, Society of Economic Paleontologists and Mineralogists Special Publication 52, p. 3–24. Kamesh Raju, K.A., Ramprasad, T., Rao, P.S., Ramalingeswara Rao, B., and Varghese, J., 2004, New insights into the tectonic evolution of the Andaman basin, northeast Indian Ocean: Earth and Planetary Science Letters v. 221, p. 145–162. Kemnitz, H., Romer, R., and Oncken, O., 2002, Gondwana break-up and the northern margin of the Saxothuringian belt (Variscides of central Europe): International Journal of Earth Sciences, v. 91, no. 2, p. 246–259, doi: 10.1007/s005310100209. Khan, P.K., and Chakraborty, P.P., 2005, Two-phase opening of Andaman Sea: A new seismotectonic insight: Earth and Planetary Science Letters, v. 229, p. 259–271, doi: 10.1016/j.epsl.2004.11.010. Klemd, R., Mathes, S., and Okrusch, M., 1991, High-pressure relics in metasediments intercaleted with the Weissenstein eclogite, Münchberg gneiss complex, Bavaria: Contributions to Mineralogy and Petrology, v. 107, p. 328–342, doi: 10.1007/BF00325102. Kossmat, F., 1927, Gliederung des varistischen Gebirgsbaues: Abhandlungen des Sächsischen Geologischen Landesamtes, v. 1, p. 1–39. Kreuzer, H., Seidl, E., Schüssler, U., Okrusch, M., Lenz, K.-L., and Raschka, H., 1989, K-Ar geochronology of different tectonic units at the northwestern margin of the Bohemian Massif: Tectonophysics, v. 157, p. 149–178, doi: 10.1016/0040-1951(89)90348-X. Krohe, A., 1996, Variscan tectonics of central Europe; postaccretionary intraplate deformation of weak continental lithosphere: Tectonics, v. 15, no. 6, p. 1364–1388, doi: 10.1029/96TC01110. Kröner, A., and Willner, A.P., 1998, Time of formation and peak of Variscan HP-HT metamorphism of quartz-feldspar rocks in the central Erzgebirge, Saxony, Germany: Contributions to Mineralogy and Petrology, v. 132, p. 1–20, doi: 10.1007/s004100050401. Kröner, A., Jäckel, P., Reischmann, T., and Kroner, U., 1998, Further evidence for an early Carboniferous (~340 Ma) age of high-grade metamorphism in the Saxothuringian Granulite complex: Geologische Rundschau, v. 86, p. 751–766. Kröner, A., Stipska, P., Schulmann, K., and Jäckel, P., 2000, Chronological constraints on the pre-Variscan evolution of the northeastern margin of the Bohemian Massif, Czech Republic, in Franke, W., Haak, V., Oncken, O., and Tanner, D., eds., Orogenic processes—Quantification and modelling in the Variscan belt of central Europe: London, Geological Society of London Special Publication 179, p. 175–197. Kroner, U., 1995, Postkollisionale Extension am Nordrand der Böhmischen Masse—Die Exhumierung des Sächsischen Granulitgebirges: Freiberger Forschungshefte, v. C, no. 457, p. 1–114. Kroner, U., and Hahn, T., 2004, Sedimentation, Deformation und Metamorphose im Saxothuringikum während der variszischen Metamorphose—Die komplexe Entwicklung Nord-Gondwanas während kontinentaler Subduktion und schiefer Kollision: Geologica Saxonica, v. 48/49, p. 133–146. Küstner, W., 2000, K-Ar-Datierungen an detritischen Muskoviten und Sm-NdModellalter prä- und synorogener schwach metamorpher Sedimente des Rhenoherzynikum—Grundlegende Daten zur Quantifizierung orogener Prozesse am Beispiel der Varisziden [Ph.D. thesis]: Göttingen, GeorgAugust Universität Göttingen, 85 p. Lange, G., Motz, H., Reichardt, C., and Schmidt, H., 1999, Stratigraphie und Tektonik nördlich und nordöstlich von Ronneburg: Beiträge zur Geologie von Thüringen, neue Folge, v. 6, p. 139–157. Langenstrassen, F., 1983, Neritic sedimentation of the Lower and Middle Devonian in the Rheinisches Schiefergebirge east of the river Rhine, in Martin H., and Eder, F.W., eds., Intracontinental fold belts: Berlin, Heidelberg, New York, Springer, p. 43–77. Laumonier, B., Autran, A., Barbey, P., Cheilletz, A., Baudin, T., Cocherie, A., and Guerrot, C., 2004, Conséquences de l’absence de socle cadomien sur l’âge et la signification des séries pre-varisques (ante-Ordovicien superieur) du sud de la France (Pyrenées, Montagne Noire): Bulletin de la Société Géologique de France, v. 175, no. 6, p. 643–655, doi: 10.2113/175.6.643.
171
Leeder, M.R., 1982, Upper Palaeozoic basins of the British Isles; Caledonide inheritance versus Hercynian plate margin processes: Journal of the Geological Society, v. 139, no. 4, p. 479–491. Linnemann, U., and Buschmann, B., 1995, Die cadomische Diskordanz im Saxothuringikum (oberkambrisch-tremadocische overlap-Sequenzen): Zeitschrift für geologische Wissenschaften, v. 23, p. 707–727. Linnemann, U., and Heuse, T., 2000, The Ordovician of the Schwarzburg anticline: Geotectonic setting, biostratigraphy and sequence stratigraphy (Saxo-Thuringian terrane): Zeitschrift der deutschen geologischen Gesellschaft, v. 151, p. 471–491. Linnemann, U., and Romer, R.L., 2002, The Cadomian orogeny in SaxoThuringia, Germany: Geochemical and Nd-Sr-Pb isotopic characterisation of marginal basins with constraints to geotectonic setting and provenance: Tectonophysics, v. 352, p. 33–64, doi: 10.1016/S0040-1951(02)00188-9. Linnemann, U., and Schauer, M., 1999, Die Entstehung der Elbezone vor dem Hintergrund der cadomischen und variszischen Geschichte des Saxothuringischen Terranes—Konsequenzen aus einer abgedeckten geologischen Karte: Zeitschrift für geologische Wissenschaften, v. 27, p. 529–561. Linnemann, U., Gehmlich, M., Tichomirowa, M., Buschmann, B., Nasdala, L., Jonas, P., Lützner, H., and Bombach, K., 2000, From Cadomian subduction to Early Palaeozoic rifting: The evolution of Saxo-Thuringia at the margin of Gondwana in the light of single zircon geochronology and basin development (central European Variscides, Germany), in Franke, W., Haak, V., Oncken, O., and Tanner, D., eds. Orogenic processes—Quantification and modelling in the Variscan belt of central Europe: London, Geological Society of London Special Publication 179, p. 131–153. Linnemann, U., McNaughton, N.J., Romer, R.L., Gehmlich, M., Drost, K., and Tonk, C., 2004, West African provenance for Saxo-Thuringia (Bohemian Massif): Did Armorica ever leave pre-Pangean Gondwana?—U/ Pb-SHRIMP zircon evidence and the Nd-isotopic record: International Journal of Earth Sciences, v. 93, p. 683–705, doi: 10.1007/s00531-0040413-8. Lorenz, W., and Hoth, K., 1990, Lithostratigraphie im Erzgebirge—Konzeption, Entwicklung, Probleme und Perspektiven: Abhandlungen des Staatlichen Museums für Mineralogie und Geologie zu Dresden, v. 37, p. 7–35. Malod, J.A., and Kemal, B.M., 1996, The Sumatra margin; oblique subduction and lateral displacement of the accretionary prism, in Hall, R., and Blundell, D.J., eds., Tectonic evolution of Southeast Asia: London, Geological Society of London Special Publication 106, p. 19–28. Mansy, J.L., Everaerts, M., and De Vos, W., 1999, Structural analysis of the adjacent Acadian and Variscan fold belts in Belgium and northern France from geophysical and geological evidence: Tectonophysics, v. 309, no. 1–4, p. 99–116, doi: 10.1016/S0040-1951(99)00134-1. Massonne, H.-J., 1999, A new occurrence of microdiamonds in quartzo-feldspatic rocks of the Saxonian Erzgebirge, Germany, and their metamorphic evolution, in Nixon, P.H., ed., Proceedings of the Seventh International Kimberlite Conference, Cape Town, April 11–17: Cape Town, Red Roof Design, p. 533–539. Mathé, G., 1969, Die Metabasite des Sächsischen Granulitgebirges: Freiberger Forschungsheft, v. C, no. 251, p. 1–130. Mathé, G., 1990, Zur Geologie der Serpentinitvorkommen im sächsischen Erzgebirge: Abhandlungen des Staatlichen Museums für Mineralogie und Geologie zu Dresden, v. 37, p. 55–72. Matte, P., 1991, Accretionary history and crustal evolution of the Variscan belt in western Europe: Tectonophysics, v. 196, p. 309–337, doi: 10.1016/00401951(91)90328-P. Mazur, S., Turniak, K., and Bröcker, M., 2004, Neoproterozoic and CambroOrdovician magmatism in the Variscan Klodzko metamorphic complex (West Sudetes, Poland): New insights from U/Pb zircon dating: International Journal of Earth Sciences, v. 93, no. 5, p. 758–772, doi: 10.1007/ s00531-004-0417-4. McCaffrey, R., 1996, Slip partitioning at convergent plate boundaries of SE Asia, in Hall, R., and Blundell, D.J., eds., Tectonic evolution of Southeast Asia: London, Geological Society of London Special Publication 106, p. 3–18. Menning, M., and German Stratigraphic Commission, 2002, A geologic time scale 2002, in German Stratigraphic Commission, eds., Stratigraphic table of Germany 2002: Potsdam. Mingram, B., 1998, The Erzgebirge Germany—A subducted part of northern Gondwana: Geochemical evidence for repetition of early Palaeozoic metasedimentary sequences in metamorphic thrust units: Geological Magazine, v. 135, p. 785–801, doi: 10.1017/S0016756898001769.
172
Kroner et al.
Mingram, B., and Rötzler, K., 1999, Geochemische, petrologische und geochronologische Untersuchungen im Erzgebirgskristallin—Rekonstruktion eines Krustenstapels: Schriftenreihe für Geowissenschaften, v. 9, p. 1–80. Mingram, B., Kröner, A., Hegner, E., and Krentz, O., 2004, Zircon ages, geochemistry, and Nd isotopic systematics of pre-Variscan orthogneisses from the Erzgebirge, Saxony (Germany), and geodynamic interpretation: International Journal of Earth Sciences, v. 93, p. 706–727, doi: 10.1007/ s00531-004-0414-7. Morley, C.K., 2004, Nested strike-slip duplexes, and other evidence for Late Cretaceous-Palaeogene transpressional tectonics before and during IndiaEurasia collision, in Thailand, Myanmar and Malaysia: Journal of the Geological Society, v. 161, no. 5, p. 799–812. Murphy, J.B., Pisarevsky, S.A., Nance, R.D., and Keppie, J.D., 2004, Neoproterozoic–Early Paleozoic evolution of peri-Gondwanan terranes: Implications for Laurentia-Gondwana connections: International Journal of Earth Sciences, v. 93, no. 5, p. 659–682, doi: 10.1007/s00531-004-0412-9. Nagy, E.A., Samson, S.D., and D’Lemos, R.S., 2002, U-Pb geochronological constraints on the timing of Brioverian sedimentation and regional deformation in the St. Brieuc region of the Neoproterozoic Cadomian orogen, northern France: Precambrian Research, v. 116, no. 1–2, p. 1–17, doi: 10.1016/S0301-9268(01)00235-2. Nance, R.D., and Murphy, J.B., 1994, Contrasting basement isotopic signatures and the palinspastic restoration of peripheral orogens: Example from the Neoproterozoic Avalonian-Cadomian belt: Geology, v. 22, p. 617–620, doi: 10.1130/0091-7613(1994)022<0617:CBISAT>2.3.CO;2. Nance, R.D., and Murphy, J.B., 1996, Basement isotopic signatures and Neoproterozoic palaeogeography of Avalonian-Cadomian and related terranes in the circum–North Atlantic, in Nance, R.D., and Thompson, M.D., eds., Avalonian and related Peri-Gondwanan terranes of the circum–North Atlantic: Boulder, Colorado, Geological Society of America Special Paper 304, p. 333–346. Nance, R.D., Murphy, J.B., and Keppie, J.D., 2002, A Cordilleran model for the evolution of Avalonia: Tectonophysics, v. 352, p. 11–31, doi: 10.1016/ S0040-1951(02)00187-7. Oncken, O., 1997, Transformation of a magmatic arc and an orogenic root during oblique collision and its consequence for the evolution of the European Variscides (Mid-German crystalline rise): Geologische Rundschau, v. 86, p. 2–20, doi: 10.1007/s005310050118. Oncken, O., 1998, Evidence for precollisional subduction erosion in ancient collisional belts: The case of the mid-European Variscides: Geology, v. 26, p. 1075–1078, doi: 10.1130/0091-7613(1998)026<1075:EFPSEI>2.3.CO;2. Pin, C., Ortega Cuesta, L.A., and Gil Ibarguchi, I., 1992, Mantle-derived, Early Palaeozoic A-type meta-granitoids from the NW Iberian Massif: Nd isotope and trace-elements constraints: Bulletin de la Société Géologique de France, v. 163, no. 4, p. 483–494. Reinhardt, J., and Kleemann, U., 1994, Extensional unroofing of granulitic lower crust and related low-pressure, high temperature metamorphism in the Saxonian Granulite Massif, Germany: Tectonophysics, v. 238, p. 71– 94, doi: 10.1016/0040-1951(94)90050-7. Robardet, M., 2002, Alternative approach to the Variscan belt in southwestern Europe: Preorogenic paleobiogeographical constraints, in Martínez Catalán, J.R., Hatcher, R.D., Jr, Arenas, R., and Díaz García, F., eds., Variscan Appalachian dynamics: The building of the Late Paleozoic basement: Boulder, Colorado, Geological Society of America Special Paper 364, p. 1–15. Robardet, M., 2003, The Armorica “microplate”: Fact or fiction? Critical review of the concept and contradictory paleobiogeographical data: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 195, p. 125–148, doi: 10.1016/S0031-0182(03)00305-5. Romer, R.L., and Rötzler, J., 2001, P-T-t evolution of ultrahigh-temperature granulites from the Saxon Granulite Massif, Germany: Part II: Geochronology: Journal of Petrology, v. 42, p. 2015–2032. Romer, R.L., and Rötzler, J., 2003, Effect of metamorphic reaction history on the U-Pb dating of titanite, in Vance, D., Müller, A., and Villa, I., eds., Geochronology: Linking the isotopic record with petrology and textures: London, Geological Society of London Special Publication 220, p. 147–158. Romer, R.L., Förster, H.-J., and Breitkreuz, Chr., 2001, Intracontinental extensional magmatism with a subduction fingerprint: the late Carboniferous Halle volcanic complex (Germany): Contributions to Mineralogy and Petrology, v. 141, p. 201–221.
Rötzler, J., and Romer, R.L., 2001, P-T-t evolution of ultrahigh-temperature granulites from the Saxon Granulite Massif, Germany: Part I: Petrology: Journal of Petrology, v. 42, p. 1995–2013. Rötzler, J., Romer, R.L., Budzinski, H., and Oberhänsli, R., 2004, Ultrahightemperature granulites from Tirschheim, Saxon Granulite Massif, Germany: P-T-t path and geotectonic implications: European Journal of Mineralogy, v. 16, p. 917–937, doi: 10.1127/0935-1221/2004/0016-0917. Russe, B., 1995, Probleme der Überschiebungstektonik im Ronneburger Erzfeld: Zeitschrift für geologische Wissenschaften, v. 23, p. 769–776. Sanchez-Garcia, T., Bellido, F., and Quesada, C., 2003, Geodynamic setting and geochemical signatures of Cambrian-Ordovician rift-related igneous rocks (Ossa-Morena Zone, SW Iberia): Tectonophysics, v. 365, no. 1–4, p. 233–255, doi: 10.1016/S0040-1951(03)00024-6. Schäfer, F., Oncken, O., Kemnitz, H., and Romer, R.L., 2000, Upper-plate deformation during collision orogeny: A case study from the German Variscides (Saxo-Thuringian zone), in Franke, W., Haak, V., Oncken, O., and Tanner, D., eds., Orogenic processes—Quantification and modelling in the Variscan belt of central Europe: London, Geological Society of London Special Publication 179, p. 281–302. Schmädicke, E., 1994, Die Eklogite des Erzgebirges: Freiberger Forschungshefte, v. C, no. 456, p. 1–344. Sherlock, S.C., Kelley, S.P., Zalasiewicz, J.A., Schofield, D.I., Evans, J.A., Merriman, R.J., and Kemp, S.J., 2003, Precise dating of low-temperature deformation; strain-fringe analysis by 40Ar/39Ar laser microprobe: Geology, v. 31, no. 3, p. 219–222, doi: 10.1130/0091-7613(2003)031<0219: PDOLTD>2.0.CO;2. Soper, N.J., and Woodcock, N.H., 2003, The lost lower Old Red Sandstone of England and Wales: A record of post-Iapetan flexure or Early Devonian transtension?: Geological Magazine, v. 140, no. 6, p. 627–647, doi: 10.1017/S0016756803008380. Stern, C.R., and Wyllie, P.J., 1981, Phase relationships of I-type granite with H2O to 35 kilobars: The Dinkey Lajes biotite-granite from the Sierra Nevada batholith: Journal of Geophysical Research, v. 86, p. 412–420. Stosch, H.G., and Lugmair, G.W., 1990, Geochemistry and evolution of MORBtype eclogites from the Münchberg Massif, southern Germany: Earth and Planetary Science Letters, v. 99, p. 230–249, doi: 10.1016/0012821X(90)90113-C. Tait, J.A., Bachtadse, V., Franke, W., and Soffel, H.C., 1997, Geodynamic evolution of the European Variscan fold belt: Paleomagnetic and geological constraints: Geologische Rundschau, v. 86, p. 585–598, doi: 10.1007/ s005310050165. Tichomirowa, M., Berger, H.-J., Koch, E.A., Belyatski, B.V., Götze, J., Kempe, U., Nasdala, L., and Schaltegger, U., 2001, Zircon ages of high-grade gneisses in the eastern Erzgebirge (central European Variscides)—Constraints on origin of the rocks and Precambrian to Ordovician magmatic events in the Variscan fold belt: Lithos, v. 56, p. 303–332, doi: 10.1016/ S0024-4937(00)00066-9. Turniak, K., Mazur, S., and Wysoczanski, R., 2000, SHRIMP zircon geochronology and geochemistry of the Orlica-Snieznik gneisses (Variscan belt of central Europe) and their tectonic implications: Geodinamica Acta, v. 13, p. 293–312, doi: 10.1016/S0985-3111(00)01045-7. von Quadt, A., 1993, The Saxonian Granulite Massif: New aspects from geochronological studies: Geologische Rundschau, v. 82, p. 516–530, doi: 10.1007/BF00212414. Werner, O., and Lippolt, H.J., 2000, White-mica 40Ar/39Ar ages of Erzgebirge metamorphic rocks: Simulating the chronological results by a model of Variscan crustal imbrication, in Franke, W., Haak, V., Oncken, O., and Tanner, D., eds., Orogenic processes—Quantification and modelling in the Variscan belt of central Europe: London, Geological Society of London Special Publication 179, p. 323–336. Wiefel, H., 1976, Die geologische Entwicklung der Lahn-Dill-Erzlagerstätte Görkwitz bei Schleiz (Oberdevon und tiefes Dinant, Thüringisches Schiefergebirge): Jahrbuch für Geologie, v. 5/6 (für 1969/70), p. 451–588. Willner, A.P., Rötzler, K., and Maresch, W.V., 1997, Pressure-temperature and fluid evolution of quartzo-feldspathic metamorphic rocks with a relic high-pressure, granulite-facies history from the central Erzgebirge (Saxony, Germany): Journal of Petrology, v. 38, p. 307–336, doi: 10.1093/ petrology/38.3.307. MANUSCRIPT ACCEPTED BY THE SOCIETY 3 OCTOBER 2006
Printed in the USA
Geological Society of America Special Paper 423 2007
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) Christoph Breitkreuz* Technische Universität Bergakademie Freiberg, D-09599 Freiberg, Germany Allen Kennedy Curtin University, Perth, Australia Marion Geißler Technische Universität Bergakademie Freiberg, D-09599 Freiberg,Germany Bodo-Carlo Ehling Saxony-Anhalt State Survey for Geology and Mining, Halle, Germany Jürgen Kopp Brandenburg State Survey for Geoscience and Resources, Kleinmachnow, Germany Andrzej Muszynski Adam Mickiewicz University, Poznan, Poland Aleksander Protas Geonafta, Piła, Poland Svend Stouge Geological Survey of Denmark and Greenland, Copenhagen, Denmark
ABSTRACT Sensitive high-resolution ion microprobe (SHRIMP) U-Pb ages have been obtained from zircons separated from Upper Carboniferous to Lower Permian SiO2rich volcanic and subvolcanic rocks of eleven drill sites. The volcanic rocks belong to a large volcanic province that formed during the initial stage of the Central European Basin System. Two drill sites are located in Denmark (North Sea and Lolland), five in northern Germany, and four in western Poland. Apart from establishing the emplacement age of the volcanic units, the focus of the present study was the dating of inherited zircons. They give information about *E-mail:
[email protected]. Breitkreuz, C., Kennedy, A., Geißler, M., Ehling, B.-C., Kopp, J., Muszynski, A., Protas, A., and Stouge, S., 2007, 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), 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. 173–190, doi: 10.1130/2007.2423(07). For permission to copy, contact
[email protected]. ©2007 Geological Society of America. All rights reserved.
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Breitkreuz et al. the chronostratigraphic structure of the anatectic component of the hybrid intraplate magmas, that is, the crust below the Central European Basin System. The Central European Basin System substrate consists mainly of a little known, covered terrane called Far Eastern Avalonia. The northern margin of the Central European Basin System rests on the collisional suture of Avalonia with Laurentia and Baltica. The southern margin is superposed on the Variscan orogen, which formed as a result of the Carboniferous collision of the Armorican Terrane Assemblage with Laurussia/ Old Red continent. Where possible, a minimum of ten analyses has been used to calculate the emplacement age for each sample, which range between 290 and 303 Ma. The two Danish samples yielded only a small zircon population with a similar, but poorly constrained emplacement age. These results, together with the data of a precursor project (Breitkreuz and Kennedy, 1999) indicate a remarkable synchronicity of the magmatic activity focused between 295 and 299 Ma throughout the Central European Basin System. Most zircon separates contain inherited grains or old cores. About seventy analyses, including data from the precursor study, are presumed to represent Avalonia and subordinately southern Baltica crust. The ages range from 320 to 2614 Ma. Paleozoic ages fall within the range of the Caledonian and Variscan orogenies. The Precambrian ages show two maxima, one between 1400 and 1600 Ma with a peak at 1450 Ma, and one between 750 and 1200 Ma with a peak at 1050 Ma. This age distribution strongly indicates an affiliation of the sub-Central European Basin System crust to Avalonia and an original position adjacent to the Amazonian craton. Keywords: SHRIMP zircon ages, inherited zircons, Precambrian, Paleozoic, central Europe, Far Eastern Avalonia, Permo-Carboniferous volcanic rocks
INTRODUCTION During the Carboniferous-Permian transition, the Central European Basin System formed (Scheck-Wenderoth and Lamarche, 2005), its southern part resting on the decaying Variscan orogen, and in the north on the Variscan foreland (Ziegler, 1990; Franke, 2000). A large part of the crustal basement below the Central European Basin System resembles Far Eastern Avalonia (Pharaoh, 1999; Banka et al., 2002), situated between Baltica (actually the northwestern part of Fennosarmatia sensu Stille 1926) in the north and the Armorican Terrane Assemblage in the south (Matte, 2001; Linnemann and Romer, 2002; Linnemann et al., 2003; see Fig. 3 below). Apart from small gneiss outcrops in the Harz Mountains (Ecker Gneiss; Baumann et al., 1991) and in the Hunsrück Mountains (Wartenstein Gneiss; Franke, 2000), both located south of the Variscan Front (SH and HM in Fig. 3), Far Eastern Avalonia basement is not exposed and its composition and age are poorly known. The magnetic and density structure of northern Far Eastern Avalonia has been revealed by geophysical experiments (Bayer et al., 2002; Williamson et al., 2002). Avalonia is supposed to have been separated from the northern margin of Gondwana during the Early Ordovician, and, drifting northward, collided with Laurentia and Baltica during the Caledonian orogeny, forming Laurussia (Old Red continent) (Pharaoh, 1999). In western England and Wales, remnants of a magmatic arc are exposed that was active on eastern Avalonia during the Ordovician (Winchester et al., 2002). Isotope studies
on eastern England and Belgian basement rocks indicate a southeastward prolongation of the arc (André et al., 1986; Noble et al., 1993). During the Variscan orogeny (Late Devonian to Carboniferous), Gondwana and the Armorican Terrane Assemblage merged with the Old Red continent, thus assembling Pangea. In a previous study, Breitkreuz and Kennedy (1999) published sensitive high-resolution ion microprobe (SHRIMP) data of old zircons that have been separated from Permo-Carboniferous volcanic rocks exposed in outcrops and drillings in northeastern Germany (Fig. 1). Based on the assumption that these inherited zircons originated from anatectic sources below the Central European Basin System, the ages revealed a complex Proterozoic and Paleozoic history of this hidden crustal segment in central Europe. The zircon separates were also used to provide for the first time reliable emplacement ages of the volcanic rocks that range between 297 and 302 Ma (with errors of ± 3 Ma) throughout the northeastern German part of the Central European Basin System. Extending this concept, we present here SHRIMP data of zircons separated from Permo-Carboniferous volcanic rocks of eleven drill cores located in the Danish North Sea and eastern Denmark, in northern Germany, and in western Poland. GEOLOGICAL SETTING OF THE DATED VOLCANIC ROCKS The dated volcanic rocks formed in a large intracontinental volcanic zone that was active during the initial phase of the Central
Jyl
lan
C-1 303
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d
Chronostratigraphic structure of Far Eastern Avalonia
Copenhagen Sjælland
55 N
Lolland Rødby 290
Daszewo 293
Fehmarn 295
Friedland 298
Hamburg Dageförde 296
Mirow 299 Penkun 300 Salzwedel 294
Kotzen 297
Flechtingen 302
Hannover
Wysoka 295
Pniewy 298
Tuchen 296
Berlin
Zdrój 296
Poznan
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Figure 1. Distribution of Permo-Carboniferous volcanic rocks (pink area, mainly subcrop) in the Central European Basin System (modified after Heeremans et al., 2004). The locations and names of the drillings are given together with the 206Pb/238U emplacement ages in Ma (see Table 1). Green circles, this study; red circles, data from Breitkreuz and Kennedy (1999).
European Basin System at the Carboniferous-Permian transition (Neumann et al., 2004; Fig. 1). The intense short-lived magmatic pulse is believed to have been controlled by dextral transtension and/or transpression (Arthaud and Matte, 1977; Lorenz and Nicholls, 1984; Ziegler, 1990). Locally, a Late Permian (Upper Rotliegend II) basaltic volcanic phase has been detected in some eastern German drill cores (Benek et al., 1996); however, this late volcanic activity is not considered here. Located 130 km west of Berlin, the Flechtingen-Roßlau uplifted block is the only outcrop of Permo-Carboniferous volcanic rocks from the Central European Basin System (Benek et al., 1973; Breitkreuz and Kennedy, 1999; Awdankiewicz et al., 2004). However, extensive hydrocarbon drilling penetrated mainly calc-alkaline volcanic rocks in a large area reaching from the southern North Sea and southern Denmark through northern Germany and into western Poland (Abragawa, 1993; Jackowicz, 1994; Marx et al., 1995; Benek et al., 1996; Maliszewska et al., 2003). SiO2-rich lava complexes and ignimbrite sheets are the predominant volcanic facies, in places piled up to 2 km or more
(Korich, 1992; Geißler and Breitkreuz, 2004; Paulick and Breitkreuz, 2005). Geochemistry, isotopic data, and the abundance of inherited zircons reveal the hybrid character of the magmas that formed from mantle melts and anatectic crustal components (Marx et al., 1995; Benek et al., 1996; Breitkreuz and Kennedy, 1999). For the first time, Late Paleozoic volcanic rocks from Danish drillings have been dated by the U-Pb method. The cores selected for the present study come from drilling C-1 located in the eastern North Sea near Jylland Peninsula. Rødby-2 is located on Lolland in the Baltic Sea (Abragawa, 1993; Figs. 1 and 2). Not far from Lolland, the drilling Fehmarn Z1 is located on the German island Fehmarn in the southwestern part of the Baltic Sea. The drilling Dageförde Z1a in eastern Lower Saxony and the drilling Salzwedel 2/64 in northern Saxony-Anhalt expose volcanic rocks of the important Altmark volcanic subprovince (Hoffmann and Stiewe, 1994; Benek et al., 1996). Located north of Berlin, the drilling Tuchen 1/74 has cores of a strongly porphyritic subvolcanic body that intruded Lower Rotliegend sediments (Gaitzsch,
3800
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Devonian limestones
dolomite and anhydrite of Lower Carboniferous
Upper Carboniferous silt- and mudstones
dacite/rhyolite lava, weakly porphyritic
293 ± 2 Ma
sediments of UR II (sandstones, minor conglomerates)
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303 ± 10 Ma
rhyolitic lava
Upper Permian (Zechstein)
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rhyolitic lava
296 ± 2 Ma
lapillistone; weakly porphyritic ?dacitic lavabreccia; tuff
sediments of UR II
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Depth CS [m]
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295 ± 2 Ma
porphyritic rhyolitic lava dome, with large feldspar phenocrysts
rhyolitic lava
sediments of UR II
Drilling Fehmarn Z1
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Drilling Dageförde Z1a
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Drilling C-1 (Danish North Sea) Profile
weakly porphyritic rhyodacitic lava (strongly flow banded)
Profile
296 ± 3 Ma
?dacite lava
sediments of UR II (sandstones, minor conglomerates)
TD: 4000,0
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Depth CS [m]
Profile
pyro-/ volcanoclastic and fluvial sediments
tuffite (?)
ignimbrite (breccia)
299 ± 3 Ma
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sediments (tuff, tuffite and fineclastic sediments) of LR
sediments of UR II
Drilling Pniewy-3
TD: 3125,0
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Depth CS [m]
290 ± 3 Ma
sediments of UR II
Drilling Zdrój-1
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Drilling Rødby-2 Profile
ignimbrite
rhyolitic tuff succession
294 ± 3 Ma
porphyritic rhyolitic lava dome
ignimbrite
sediments of UR II
TD: 4326,5
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Depth CS [m]
Profile
296 ± 3 Ma
intrusive, porphyritic rhyolite
sediments of LR (+ ?UR I)
sediments of UR II
Drilling Tuchen 1/74
TD: 4987,85
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Drilling Salzwedel 2/64
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298 ± 3 Ma, (Breitkreuz and Kennedy 1999)
moderately plag-phyric dacite lava
aphyric andesite intrusion
moderately plag-phyric dacite lava
moderately plag-Kfsp-phyric dacite lava intercalated with thinnly bedded, normal graded, tuffaceous mud and sandstone
300 ± 5 Ma
highly plag-Kfsp-phyric dacite lava
sediments of UR II
Drilling Penkun 1/71
Breitkreuz et al.
Figure 2. Schematic drilling sections showing the segments from which the samples have been taken (1–2 m of half cores for each sample) and the emplacement age (Fig. 1). Ages have been rounded; for exact values see Table 1 and the Appendix. Drillings are displayed in different scales. cs—core segment; Kfsp—K-feldspar; LR—Lower Rotliegend (Latest Carboniferous to Lower Permian); plag—plagioclase; TD—total depth; UR I—Upper Rotliegend I; UR II—Upper Rotliegend II (Late Permian).
Devonian limestones
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TD: 4000
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TD: 3873
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porphyritic rhyolite lava with feldspar phenocrysts (< 0.5 cm); 1m tuffaceous conglomerate at the top
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295 ± 2 Ma
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sediments of UR II (fine grained sandstones with mudstones); at the bottom conglomerate with clasts up to 5 mm
Profile
Depth [m] CS
Drilling Daszewo-12
Profile
Depth [m]
CS
Drilling Wysoka Kamienska-2
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Chronostratigraphic structure of Far Eastern Avalonia 1995). Finally, the uppermost unit of the 1500-m-thick volcanic succession exposed in the drilling Penkun 1/71 has been included in the present project. The lowermost unit was already dated by Breitkreuz and Kennedy (1999; Fig. 2). We also present the first U-Pb ages of Polish Permo-Carboniferous volcanic rocks exposed in drillings. Samples come from drilling Daszewo-12 in Pomerania, from drilling Wysoka Kamienska-2 near Szczecin close to the Baltic Sea, and from drillings Pniewy-3 and Zdrój-1 near Poznan.
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We carried out the same mounting and analytical procedures at the SHRIMP in the Curtin University in Perth that were applied in the precursor project (description in Breitkreuz and Kennedy, 1999). Where possible, we used ten or more spot measurements on different zircon grains to calculate the 206Pb/238U emplacement age of the volcanic unit (Table 1 and Appendix). For zircons older than 800 Ma we have used the 207Pb/206Pb age. The 207Pb/206Pb age has also been adopted if the analysis of a younger zircon is highly discordant.
METHODS RESULTS Ten samples have been taken from lithic-free dacitic to rhyolitic lavas or subvolcanic intrusions (Fig. 2). The sample of drilling Pniewy-3 has been taken from an ignimbrite. The lithic clasts in this ignimbrite have been handpicked after crushing the rock to pea-sized fragments to minimize contamination by zircons that were not located in the magma chamber prior to eruption. Silica-rich porphyritic volcanic rocks (lava, ignimbrite, and subvolcanic rocks) are normally rich in zircons and therefore suitable for this kind of project. With the exception of the weakly porphyritic samples of the Danish cores (C-1 and Rødby-2) and of the Polish sample from Pniewy-3, the selected samples yielded abundant zircons.
The complete data set of SHRIMP measurements is listed in the Appendix; a summary is given in Table 1. The 206Pb/238U ages (bold in Table 1) are presumed to represent the emplacement age of the volcanic and subvolcanic units (Figs. 1 and 2). Because of few measurements, the accuracy of the C-1, Rødby-2, and Pniewy-3 ages is relatively poor; all other samples yielded wellcontrolled emplacement ages. This accuracy is supported by the similarity of the respective 206Pb/238U, 207Pb/235U, and 207Pb/206Pb ages (Table 1). The emplacement ages of the eleven samples range from 290 to 303 Ma and complement the seven ages from the northeastern German part of the Central European Basin
TABLE 1. SUMMARY OF ZIRCON SHRIMP DATING OF CORE SEGMENTS DEPICTED IN FIGURE 2 206 238 † 207 235 207 206 § Pb/ U age Pb/ U age Pb/ Pb age Age of inherited cores Number of (Ma) (Ma) (Ma) (Ma) zircons* C-1 2 294.2 ± 18.3 232 ± 130 1558 302.6 ± 9.6 (4 measurements) §§ Rødby-2 1 293.6 ± 3.1 323 ± 16 547, 552 289.8 ± 2.6 Dageförde Z1a 11 295 ± 4 291 ± 18 not found 296 ± 2 # # Fehmarn Z1 17 296.7 ± 4.2 306 ± 30 325**, 499 , 1078, 1502 295.3 ± 2.3 # †† Salzwedel 2/64 12 294.5 ± 5.0 294 ± 35 336 , 792**, 982, 1428 , 293.8 ± 2.7 2104 Penkun 1/71 15 299.4 ± 6.3 304 ± 32 354, 642, 1464, 1665, 299.7 ± 5.3 1793 Tuchen 1/74 10 298 ± 6 296 ± 32 323, 360, 372, 408, 447, 296 ± 3 558, 636, 722, 758, 773, 823, 878, 892, 967, 969, 976, 1123, 1238, 1255, 1354, 1393, 1394, 1415, 1456, 1465, 1618, 1660, 1831, 1873, 2245, 2614 Zdrój-1 12 292 ± 5 273 ± 42 338, 602, 1124, 1492 296 ± 3 §§ 299.4 ± 4.7 305 ± 27 599, 984, 1020 Pniewy-3 5 298.5 ± 3.3 (7 measurements) Wysoka Kamienska-2 18 292.9 ± 3.9 283 ± 28 not found 294.6 ± 2.3 Daszewo-12 21 292.7 ± 4.2 293 ± 30 726, 1260, 2094, 2344 293.0 ± 2.3 Notes: Assumed emplacement ages are listed in bold. Complete data set is given in the Appendix. Error is 2σ. *Used for the calculation of the emplacement age. † Assumed emplacement age of the volcanic unit. § 206 328 207 206 Below 500 Ma the Pb/ U age, above 500 Ma the Pb/ Pb age is given; for exceptions, see footnotes below. # 206 238 Pb/ U data from two grains used for calculation. 206 238 ** Pb/ U data from three grains used for calculation. †† 206 238 Pb/ U data from four grains used for calculation. §§ Unrealistically small errors of the Rødby-2 and Pniewy-2 emplacement ages, because of the small number of measurements. Drilling
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C-1 1558
Rødby 547, 552
Daszewo 726, 1260, 2094, 2344
Fehmarn 325, 499, 1078, 1502
Friedland 387, 443,1456 Salzwedel 336, 792, 982, 1428, 2104 Dageförde Flechtingen 350, 538
Wysoka
Penkun 354, 642, 1665, 1464, 1483 1793 Pniewy Tuchen 599, 984, 1020 323 - 2614, n = 31 Zdrój Kotzen 338, 602, 1124, 1492, 1544 345 Halle Volc. Complex 325-398,500-650,1674, 1843, 2065-2373, n = 14
Figure 3. Map of the Paleozoic terrane assemblage of central Europe (Banka et al., 2002). For a key to drillings see Figure 1. Numbers are ages of inherited zircons (207Pb/206Pb ages are shown for Precambrian zircons; see Appendix), including data from Breitkreuz and Kennedy (1999). CDF—Caledonian deformation front; EL—Elbe lineament; HM—Hunsrück Mountains; SH—South Hunsrück.
System presented by Breitkreuz and Kennedy (1999). The frequency distribution of all eighteen ages has a maximum between 295 and 299 Ma. The sample locations spread over an area of ~800 × 500 km2; however, no regional trend is apparent (Fig. 1). Ages of inherited zircons or of old zircon cores are summarized in the right column of Table 1 and displayed in concordia plots grouped according to the drilling locations (Fig. 4). Furthermore, a map view (Fig. 3) and a histogram (Fig. 5A) show the results together with ages determined in the previous study (Breitkreuz and Kennedy, 1999). The samples from Dageförde Z1a and Wysoka Kamienska-2 yielded good emplacement ages (Fig. 2, Table 1); however, no inherited zircons have been detected. The same observation was made by Breitkreuz and Kennedy (1999) for two samples of the drilling Mirow 1/74 and one sample of Friedland 1/71. The lack of inherited grains in
these samples may stem from a number of reasons: (1) the volcanic rocks have a zirconium-poor anatectic and/or mantle source; or (2) they formed from superheated melt, which annealed any inherited component. All other mineral separates yielded old zircon ages, more than sixty-three in total. Paleozoic ages cluster in the time spans 360–325 and 499–400 Ma. The Precambrian ages range from 560 to 2614 Ma, thus spanning the entire Proterozoic and the Late Archean. DISCUSSION Emplacement Ages of the Volcanic Rocks The 206Pb/238U emplacement ages presented here complement the ages published by Breitkreuz and Kennedy (1999;
0
0
400
600
C
206Pb 238U
400
600
A
1
800
1
800
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1400
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4
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Rødby-2
Fehmarn Z1
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C-1
Daszewo-12
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207Pb 235U
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206Pb 238U
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206Pb 238U
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*
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207Pb 235U
Zdrój-1
Pniewy-3
7
8
207Pb 235U
Penkun 1/71 Breitkreuz and Kennedy 1999 Tuchen 1/74
2000
6
2000
Figure 4. Concordia diagrams of inherited ages. (A) Zircons presumably originating from Baltica crust; (B) assumed Far Eastern Avalonia zircons from western Poland; (C) assumed Far Eastern Avalonia zircons from the central section of north Germany and southern Denmark; (D) assumed Far Eastern Avalonia zircons from northeastern Germany (the Late Archean age of Tuchen 1/74 is not displayed, see Table 1 and Appendix).
0.06
0.10
0.14
0.18
0.22
0.26
0.30
0.34
0.38
0.06
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206Pb 238U
Chronostratigraphic structure of Far Eastern Avalonia
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A
7 6 Baltica Avalonia
Frequency
5 4
S Fennoscandia*
3 2 1
B
2700
2500
2300
2100
1900
1700
1500
1300
1100
900
Age in Ma
Leonian
Eburnian
Pan-African
West African Craton Liberian
700
500
0
Central-Amazonian
Amazonian Craton
Trans-Amazonian
Rio Negro
Rondonian
Sunsas
Brasiliano
C
Figure 5. Distribution of Precambrian 207Pb/206Pb age of inherited zircons from our project compared to the chronostratigraphic structure of: (A) southern Baltica (* indicates data from Obst et al., 2004); (B) the West African craton; and (C) the Amazonian craton. (B and C taken from Nance and Murphy, 1994). The data shown in (A) are from twelve different samples, including data presented by Breitkreuz and Kennedy (1999). They have not been filtered based on discordance. Vertical dashed lines mark 500-Ma intervals.
Chronostratigraphic structure of Far Eastern Avalonia Fig. 1). Thus, the concept of a short, intense “magmatic flare-up” focused between 295 and 299 Ma, presumably caused by decompressional melting of the upper mantle, might be extended to the entire Central European Basin System volcanic zone. Within the large area considered, no regional trend is apparent, an observation that renders an alternative model of an intracontinental hot spot trace unlikely (Smith and Braile, 1994). Numerous studies have attempted K/Ar dating of Central European Basin System volcanic rocks, yielding mostly Triassic to Jurassic “ages.” The disturbance of the K/Ar system is presumably related to a strong Jurassic thermal event, which affected large areas of central Europe (Zwingmann et al., 1998; Brecht, 1999). From the Poznan area, an Ar/Ar age of 288 ± 5 Ma (Lippolt et al., 1982) and a K/Ar age of 279 ± 7 Ma (Breitkreuz et al., 2000) have been reported. The biotite separate for the latter was gained from the lower core segment of Pniewy-3 (see Fig. 2). Our SHRIMP data of Zdrój-1 and Pniewy-3 (Table 1) reveal that the Ar/Ar and K/Ar data of western Poland are also not reliable. Likewise, various K/Ar datings of volcanic rocks from Danish drill cores yielded presumably disturbed Mid- to Late Permian ages (reviews in Stemmerik et al., 2000; Heeremans and Faleide, 2004). Origin of Paligenic Zircons Magmatogenetic concepts of Permo-Carboniferous volcanism in the Central European Basin System and of the possible origin of paligenic zircons have been discussed by Breitkreuz and Kennedy (1999). It is likely that the voluminous calc-alkaline volcanics were generated in a relatively short time from hybrid magmas that formed as a consequence of the intrusion of mantle melts into the lower crust below the Central European Basin System (see also Romer et al., 2001). Although most of the palingenic zircons are assumed to have originated from those parts of the lower crust that experienced anatexis, minor assimilation of middle to upper crustal material during ascent and differentiation of the magmas cannot be ruled out. Considering the limited number of measurements, the age spectra displayed in Figure 4 are quite variable. A dense cluster of ages (as, e.g., apparent from the Penkun and Salzwedel zircons) presumably reflects a (meta)magmatic or first-cycle metasedimentary source (Fig. 4C, D). In contrast, the Tuchen zircon population strikes by its wide and continuous age spectrum (Fig. 4D), which might represent an additional multicycle metasedimentary source. Paleozoic Inherited Ages Considering the Paleozoic plate assemblage of Europe outlined above (Fig. 3), remnants of an Ordovician/Silurian magmatic arc should exist in Far Eastern Avalonia because the suture between Baltica and Far Eastern Avalonia is inclined toward the south (Meissner and Krawczyk, 1999; Bayer et al., 2002), and because there is no trace of Ordovician/Silurian magmatism on
181
Baltica east of the Norwegian Caledonides (Obst et al., 2004). A “lost magmatic arc” located on Far Eastern Avalonia has been postulated on the basis of sedimentological and geophysical studies (McCann, 1998; Williamson et al., 2002). The 447-Ma age from Tuchen 1/74 (Table 1) and that of 443 Ma from Friedland 1/71 (Breitkreuz and Kennedy, 1999) might represent ages of Caledonian arc magmatism in Far Eastern Avalonia. Post-Caledonian, early to late Variscan ages occur throughout much of the Central European Basin System (408–323 Ma; Breitkreuz and Kennedy, 1999; Table 1). Middle Devonian (Givet, ca. 385 Ma) volcanic activity is well known from the southern margin of the Old Red continent, that is, in the Lahn-Dill synclines of the Rhenish Slate Mountains (Nesbor et al., 1993). Late Devonian to Carboniferous magmatism also took place in Far Eastern Avalonia north of the Variscan deformation front (Fig. 3), an observation that has also been made for the British Isles and for northwestern Poland (compilation in Ziegler, 1990). Inherited Precambrian Ages Southern Baltica is characterized by a Svecofennian (1.9– 1.86 Ga) and a Sveconorwegian domain with Grenvillian overprint (1.0 and 0.9 Ga), the former domains being separated by the Transscandinavian igneous belt (1.81–1.66 Ga) (Obst et al., 2004; Fig. 5A). C-1 and Daszewo-12 drillings are clearly located on Baltica. The 1558-Ma age of C-1 belongs to the Sveconorwegian domain (i.e., Southwest Scandinavian domain; 1.8–0.9 Ga according to Gáal and Gorbatschev, 1987). The ages of 2094 and 2344 Ma of Daszewo-12 seem to be too old for the Svecofennian domain and maybe more related to the Karelide domain in eastern Baltica (2.5–1.9 Ga; Gáal and Gorbatschev, 1987). The Rødby-2 zircon population yielded, among others, ages of 547 and 552 Ma (Fig. 4C). Located <30 km southwest of Rødby, the drilling Fehmarn Z1 revealed a 499-Ma zircon, which may belong to the same chronostratigraphic unit that was sampled by the Rødby magma. Rødby-2 is located just south of the Caledonian deformation front (CDF in Fig. 3) running between the Danish islands of Lolland and Sjælland (Fig. 1). The basement block to the north of the Caledonian deformation front, (i.e., the Sveconorwegian domain) experienced the last strong metamorphic overprint as early as 1.0–0.9 Ga (Obst et al., 2004). Ages of 550–500 Ma are typical Panafrican or Brazilian ages (Nance and Murphy, 1994; Fig. 5B, C). However, latest Ediacaran ages have been reported from the Holy Cross Mountains in eastern Poland, an area for which a close proximity to Baltica has been assumed during that period (see overview in Banka et al., 2002). Thus, the terrane affiliation (Baltica versus Far Eastern Avalonia) of the latest Precambrian to Cambrian sources untapped by the Rødby and Fehmarn magmas remains unclear. The supracrustal Rheno-Hercynian thrust and fold belt (Fig. 3) consists predominantly of sedimentary and volcanic rocks deposited at the southern continental margin of Laurussia (Old Red continent); however, in its southern part, some allochthonous nappes from the Armorican Terrane Assemblage and
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Breitkreuz et al.
syn-orogenic detritus have been transported to the north during the Variscan orogeny (Huckriede et al., 2004). The boundary between Far Eastern Avalonia and Armorican Terrane Assemblage crust is marked by the Hunsrück Taunus Quartzite and Wippra Phyllite zone at the southern margin of the RhenoHercynian orogenic unit (Franke, 2000). The spectrum of inherited Precambrian ages derived from our studies clearly differentiates Far Eastern Avalonia from Armorican Terrane Assemblage crust. Armorican Terrane Assemblage formed part of the West African craton throughout the Proterozoic, and its zircons display a prominent age gap between 1900 and 700 Ma (Nance and Murphy, 1994; Linnemann et al., 2003; Fig. 5B). For example, inherited zircons of the Halle volcanic complex located on Armorican Terrane Assemblage crust show a gap between 1674 and 650 Ma (Breitkreuz and Kennedy, 1999; Fig. 3). In contrast, the age spectrum of inherited zircons of our study shows two frequency maxima at ca. 1000 and 1500 Ma, typical for the chronostratigraphic evolution of the Amazonia craton (Fig. 5C). Bayer et al. (2002) discussed the possibility that Baltic crust reaches below the Central European Basin System as far as the Elbe lineament (EL in Fig. 3). Alternatively, according to these authors, high seismic velocities in the crust below the Central European Basin System suggest the presence of oceanic crust trapped during the Caledonian orogeny. The latter model seems unlikely, considering the abundance of Late Archean to Proterozoic zircons found in our study. McNamara et al. (2001) questioned the proximity of West Avalonia to Amazonia during the latest Precambrian on the basis of a paleomagnetic study. In contrast, Murphy et al. (2002) point to the presence of Grenvillian-age zircons in West Avalonia rocks, indicating the proximity of Avalonia to an eroding Grenvillian orogen. From the British segment of Avalonia (East Avalonia; Fig. 3) only Neoproterozoic intrusion ages are known; however, upper intercepts of discordant ages point to the existence of a MidProterozoic core (Tucker and Pharaoh, 1991). The age ranges of West Avalonia (520–500 Ma, 3000–900 Ma, with an apparent gap between 2400 and 2100 Ma; Murphy et al., 1999) compare quite well with the data yielded by the Far Eastern Avalonia zircons. However, the ages of 2104 and 2245 Ma (Salzwedel 2/64 and Tuchen 1/74, respectively) lie within this West Avalonian age gap. For the first time, a Late Archean age has been detected in a
Far Eastern Avalonia zircon (Tuchen 1/74), which is synchronous with the central Amazonian phase (Fig. 5). CONCLUSIONS SHRIMP dating of eleven zircon separates from drill cores from Denmark, Germany, and Poland reveal reliable emplacement ages for the host volcanic rocks and abundant inherited ages representing the anatectic component of the hybrid magmas. Including the data of the precursor study (Breitkreuz and Kennedy, 1999), eighteen emplacement ages of SiO2-rich volcanic rocks from a 800 × 500-km2 area of the Central European Basin System range between 290 and 303 Ma, with a focus between 295 and 299 Ma. These data include the first Late Paleozoic U/Pb datings of volcanic rocks from Denmark and Poland. There is no apparent regional trend. Fifty Precambrian and thirteen Paleozoic ages display a fairly substantial chronostratigraphy of Far Eastern Avalonia and its northern border with Baltica. However, apart from its obvious continental nature, a further lithological or geochemical characterization of the hidden Far Eastern Avalonia crust is not possible. The Paleozoic ages of Far Eastern Avalonia complement the Caledonian dynamics (amalgamation of Avalonia with Laurentia and Baltica) and the Early to Late Variscan magmatic events that occurred in the Rheno-Hercynian zone and the foreland to the north. The Precambrian age range of Far Eastern Avalonia zircons suggests a close affinity to the Amazonian craton. In particular, the two age maxima at 1050 and 1450 Ma allow for a clear distinction from the Armorican Terrane Assemblage, which displays a long age gap between 800 and 1900 Ma. ACKNOWLEDGMENTS We thank Biuro Geologiczne Geonafta; Deutsche Wissenschaftliche Gesellschaft für Erdöl, Erdgas, und Kohle; ErdgasErdöl Berlin GmbH; Geological Survey of Denmark and Greenland; and RWE DEA for permission to access and sample core material. The project has been funded by the German Research Foundation (DFG grant Br 997/19-1). Reviews by Tim Pharaoh and Wolfgang Franke helped considerably to improve and balance the manuscript.
197 248 211 291 167 167 127 149 130 294 870 156 176 96 122 180 202 137
102 151 131 189 78 79 57 73 68 256 42 131 81 25 59 82 114 71
76 120 307 654 76 105 145 840 1733 128 88 97 56
Dageförde Z1a 1 179 2 224 3 493 4 905 5 167 6 253 6b 243 7 995 8 1766 9 318 10 183 11 206 12 238
Daszewo-12 dasz-1 dasz-1b dasz-2 dasz-3 dasz-3b dasz-4 dasz-5 dasz-6 dasz-7 dasz-8 dasz-8b dasz-9 dasz-10 dasz-11 dasz-12 dasz-13 dasz-14 dasz-15
187 1828 4420 63 62
296 580 1056 182 212
C-1 c1-1 c1-1a c1-1b c1-2a c1-2b
Th (ppm)
U (ppm)
Drilling spot
APPENDIX
11 13 11 20 8 9 7 8 7 112 76 75 8 22 7 9 10 7
9 11 26 46 7 12 13 54 94 17 9 10 11
88 50 105 9 10
Pb (ppm)
2.54 2.75 2.56 8.34 0.99 3.10 2.79 0.57 0.54 2.01 0.20 0.69 0.07 0.19 6.16 0.02 0.54 0.32
0.0068 0.31 0.14 0.22 0.1 0.2 2.5 0.49 0.38 3 0.25 1.5 0.21
0.03 0.72 0.68 0.06 0.34
f (%)*
0.0473 0.0467 0.0466 0.0491 0.0469 0.0472 0.0478 0.0478 0.0482 0.3022 0.0936 0.3940 0.0466 0.2239 0.0461 0.0464 0.0457 0.0471
0.0478 0.0477 0.048 0.0461 0.0418 0.047 0.0463 0.0468 0.0449 0.0469 0.0465 0.0465 0.0468
0.2679 0.0479 0.0484 0.0485 0.0481
238
Pb/ U
206
9 10 9 10 9 9 10 10 10 28 8 42 4 24 5 4 4 5
45 44 39 36 48 45 44 37 34 42 48 47 45
24 9 10 10 9
±
0.3616 0.3252 0.3330 0.4036 0.3411 0.3410 0.3540 0.3215 0.3371 5.4084 0.8201 8.1391 0.3307 2.5506 0.3650 0.3451 0.3091 0.3158
0.3447 0.341 0.347 0.3258 0.3172 0.3357 0.3487 0.3275 0.3146 0.344 0.3367 0.3325 0.3286
3.5654 0.3031 0.3110 0.3452 0.3363
235
Pb/ U
207
143 163 146 227 140 165 167 151 153 756 90 1174 97 405 197 98 104 116
915 904 688 512 1087 911 1158 665 525 1296 1080 1126 841
429 309 382 121 112
±
0.0554 0.0505 0.0519 0.0596 0.0528 0.0524 0.0537 0.0487 0.0507 0.1298 0.0635 0.1498 0.0515 0.0826 0.0575 0.0540 0.0491 0.0486
0.0523 0.0518 0.0524 0.0513 0.055 0.0518 0.0546 0.0508 0.0508 0.0532 0.0525 0.0519 0.051
0.0965 0.0459 0.0466 0.0516 0.0507
206
Pb/ Pb
207
17 21 18 29 17 21 21 19 19 11 3 12 13 8 29 13 15 16
124 123 90 65 170 125 168 90 71 189 153 161 115
6 44 55 13 12
±
238
298 294 293 309 295 297 301 301 303 1702 577 2141 293 1302 290 292 288 297
301 301 302 290 264 296 292 295 283 295 293 293 295
1530 302 304 305 303
206
Pb/ U age (Ma)
6 6 6 6 6 6 6 6 6 14 5 20 3 13 3 3 3 3
3 3 2 2 3 3 3 2 2 3 3 3 3
13 6 6 6 6
± (Ma)
TABLE A1. COMPLETE DATA SET OF SHRIMP ANALYSES
313 286 292 344 298 298 308 283 295 1886 608 2247 290 1287 316 301 273 279
301 298 302 286 280 294 304 288 278 300 295 291 289
1542 269 275 301 294
207
235
Pb/ U age (Ma)
11 12 11 16 11 13 13 12 12 12 5 13 7 12 15 7 8 9
7 7 5 4 8 7 9 5 4 10 8 9 6
10 24 30 9 9
± (Ma)
430 219 280 591 320 303 361 136 228 2095 726 2344 264 1260 509 370 152 129
301 277 304 253 413 276 396 231 231 339 309 279 240
1558 0 40 268 229
206
Pb/ Pb age (Ma)
207
70 98 83 108 75 94 90 89 87 16 11 14 60 20 113 57 71 77
54 54 39 29 69 55 69 41 32 80 66 71 52
12 21 253 60 55
± (Ma)
69 134 105 52 92 98 83 222 133 81 79 91 111 103 57 79 190 230 Continued
100 109 100 115 64 107 74 128 122 87 95 105 123
98 0 755 114 132
Concordance (%)
Chronostratigraphic structure of Far Eastern Avalonia 183
94 134 95 108 82 41 133 121 152 164 53 161 84 90 63 231 145 39 66 44 17 15 47 114 63
170 225 234 275 139 291 325 2097 533 480 289 417 275 1392 178 452 256 122 54 120 84 64 75 185 195
Fehmarn Z1 feh-1 feh-2 feh-3 feh-4 feh-5 feh-6a feh-6b feh-6c feh-7 feh-8 feh-9 feh-10 feh-11a feh-11b feh-12a feh-12b feh-13 feh-14 feh-15 feh-16 feh-17 feh-17b feh-18 feh-19 feh-20
Th (ppm)
80 89 114 102 116 75 133 54
U (ppm)
Daszewo-12 (continued) dasz-16 150 dasz-17 176 dasz-18 186 dasz-19 169 dasz-20 251 dasz-21 147 dasz-22 223 dasz-23 115
Drilling spot
8 11 11 13 7 13 17 102 25 93 13 24 12 65 8 22 13 6 18 6 6 5 21 9 9
9 9 11 9 12 7 11 6
Pb (ppm)
0.00 0.08 0.34 0.12 0.19 0.02 0.07 0.13 0.10 0.28 0.47 6.13 0.15 0.00 0.21 0.20 0.00 0.05 0.00 0.19 0.00 0.12 0.07 0.16 0.25
7.82 0.00 5.37 3.24 0.05 0.00 0.04 0.07
f (%)*
238
0.0471 0.0468 0.0460 0.0470 0.0469 0.0471 0.0507 0.0526 0.0474 0.1892 0.0465 0.0460 0.0458 0.0503 0.0466 0.0462 0.0465 0.0459 0.2647 0.0472 0.0781 0.0820 0.2550 0.0464 0.0468
0.0463 0.0466 0.0468 0.0453 0.0465 0.0458 0.0466 0.0466 5 5 5 5 6 5 5 4 4 18 5 4 4 4 5 4 5 6 40 7 10 22 37 5 5
5 4 5 6 4 5 4 5
±
0.3480 0.3508 0.3189 0.3449 0.3401 0.3417 0.3748 0.3864 0.3346 1.9653 0.3364 0.3370 0.3299 0.3669 0.3330 0.3311 0.3451 0.3350 3.2235 0.3264 0.6250 0.6568 3.1976 0.3305 0.3408
0.3272 0.3425 0.3348 0.3289 0.3312 0.3302 0.3381 0.3288
235
Pb/ U
207
86 134 108 102 162 77 99 43 71 259 86 129 90 44 119 90 77 133 696 149 163 304 977 143 114
159 64 240 171 84 69 102 123
±
0.0535 0.0543 0.0503 0.0532 0.0526 0.0527 0.0536 0.0532 0.0512 0.0753 0.0524 0.0531 0.0523 0.0529 0.0519 0.0519 0.0539 0.0529 0.0883 0.0501 0.0580 0.0581 0.0910 0.0517 0.0528
0.0513 0.0533 0.0519 0.0526 0.0517 0.0523 0.0526 0.0511
206
Pb/ Pb
207
10 18 15 14 23 9 12 3 9 5 11 18 12 3 16 12 9 19 11 20 11 19 22 20 15
23 7 36 25 11 8 14 17
±
291 294 295 286 293 288 294 294
238
297 295 290 296 295 296 319 331 298 1117 293 290 288 317 293 291 293 290 1514 298 485 508 1464 292 295
206
Pb/ U age (Ma)
4 3 3 3 4 3 3 3 3 10 3 3 3 3 3 3 3 4 21 4 6 14 19 3 3
3 3 3 4 3 3 3 3
± (Ma)
TABLE A1. COMPLETE DATA SET OF SHRIMP ANALYSES (continued) Pb/ U
206
303 305 281 301 297 298 323 332 293 1104 294 295 289 317 292 290 301 293 1463 287 493 513 1457 290 298
287 299 293 289 290 290 296 289
207
235
Pb/ U age (Ma)
7 10 8 8 12 6 7 3 5 9 7 10 7 3 9 7 6 10 17 11 10 19 24 11 9
12 5 18 13 6 5 8 9
± (Ma)
352 385 207 338 311 314 354 339 251 1078 304 333 298 323 279 283 365 324 1390 201 530 534 1446 272 321
255 343 281 313 271 300 312 247
206
Pb/ Pb age (Ma)
207
46 78 71 60 101 42 52 14 41 16 50 81 54 15 74 55 40 82 26 95 45 74 48 92 68
107 32 159 110 51 37 62 79
± (Ma)
84 77 140 88 95 94 90 98 119 104 96 87 97 98 105 103 80 89 109 148 91 95 101 107 92 Continued
114 86 105 91 108 96 94 119
Concordance (%)
184 Breitkreuz et al.
25 47 267 259 325 128 138 204 339 189
16 3058 25
1144 678 1688
74 78 43 56 17 80 59 66 105 55 38 140 431 190 183 87 88 97 76 97 277
Th (ppm)
189 75 804 748 907 338 209 518 727 498
155 153 105 158 31 152 195 191 136 139 140 170 762 243 372 186 168 121 151 200 401
Penkun 1/71 penk-1 penk-2 penk-3 penk-4 penk-5 penk-6a penk-6b penk-7a penk-7b penk-8 penk-9 penk-9b penk-10 penk-11 penk-12 penk-13 penk-14 penk-15 penk-16 penk-17 penk-18
Pniewy-3 pni-1 pni-1b pni-2 pni-3 pni-4 pni-5 pni-6a pni-6b pni-2b pni-6c Rødby-2 rod-1 rod-2 rod-3
U (ppm)
Drilling spot
88 66 138
15 21 38 36 44 18 42 25 35 24
8 8 5 8 2 57 9 9 37 7 6 43 39 16 20 9 8 13 8 10 21
Pb (ppm)
0.04 0.00 0.08
0.06 3.32 0.00 0.30 1.11 2.93 0.25 0.00 0.00 0.03
0.37 0.00 0.00 0.44 0.16 0.00 0.00 0.07 0.00 0.13 0.21 0.00 0.34 0.32 2.20 0.01 0.30 0.00 0.20 0.05 0.03
f (%)*
238
0.0837 0.0460 0.0895
0.0859 0.2338 0.0472 0.0474 0.0464 0.0466 0.1803 0.0473 0.0464 0.0468
0.0496 0.0478 0.0473 0.0482 0.0484 0.3439 0.0487 0.0483 0.2422 0.0474 0.0468 0.2195 0.0475 0.0567 0.0479 0.0469 0.0476 0.0920 0.0476 0.0474 0.0470
7 4 8
9 31 4 4 4 4 21 4 4 4
12 10 10 10 12 72 10 10 51 10 10 46 9 11 9 10 10 19 10 9 9
±
0.6753 0.3353 0.7229
0.7089 2.3183 0.3411 0.3405 0.3255 0.3492 1.8202 0.3427 0.3383 0.3396
0.3517 0.3465 0.3668 0.3229 0.3370 5.1979 0.3601 0.3431 3.0676 0.3420 0.3285 3.0941 0.3381 0.4038 0.3497 0.3531 0.3236 0.7745 0.3384 0.3362 0.3427
235
Pb/ U
207
68 40 78
146 816 47 63 66 123 423 54 47 76
173 103 121 136 300 1213 99 122 721 136 124 771 97 163 131 133 135 202 146 125 112
±
0.0585 0.0529 0.0586
0.0599 0.0719 0.0524 0.0521 0.0509 0.0543 0.0732 0.0525 0.0529 0.0526
0.0515 0.0526 0.0562 0.0486 0.0505 0.1096 0.0536 0.0516 0.0919 0.0523 0.0509 0.1022 0.0516 0.0517 0.0530 0.0546 0.0493 0.0611 0.0516 0.0514 0.0529
206
Pb/ Pb
207
2 3 2
9 22 4 7 8 17 13 5 4 10
20 9 12 16 41 8 8 13 6 16 14 11 9 16 15 15 16 8 17 14 12
±
238
518 290 552
531 1354 297 299 292 294 1069 298 292 295
312 301 298 303 305 1905 307 304 1398 298 295 1279 299 355 301 296 300 567 300 298 296
206
Pb/ U age (Ma)
4 3 5
6 17 3 3 3 3 12 3 3 3
7 6 7 7 8 35 6 6 27 6 6 25 6 7 6 6 6 11 6 6 6
± (Ma)
TABLE A1. COMPLETE DATA SET OF SHRIMP ANALYSES (continued) Pb/ U
206
524 294 552
544 1218 298 298 286 304 1053 299 296 297
306 302 317 284 295 1852 312 300 1425 299 288 1431 296 344 304 307 285 582 296 294 299
207
235
Pb/ U age (Ma)
4 3 5
9 25 4 5 5 9 15 4 4 6
13 8 9 10 23 20 7 9 18 10 10 19 7 12 10 10 10 12 11 10 8
± (Ma)
548 324 552
599 984 303 288 236 385 1020 308 325 313
262 313 462 129 218 1793 355 266 1464 301 235 1665 267 271 329 394 164 642 267 261 323
206
Pb/ Pb age (Ma)
207
9 16 10
35 63 21 35 39 73 38 26 21 43
92 42 48 76 183 14 36 60 14 71 66 20 41 74 65 64 78 28 79 65 53
± (Ma)
95 89 100 Continued
89 138 98 104 124 76 105 97 90 94
119 96 65 234 140 106 86 114 95 99 126 77 112 131 92 75 183 88 112 114 92
Concordance (%)
Chronostratigraphic structure of Far Eastern Avalonia 185
218 28 162 125 65 190 28 47 57 608 60 721
672 560 312 438 604 774 351 45 727 2435 599 1035
Tuchen 1/74 A.1 core A.2 core A.3 interior A.4 interior A.5 interior A.6 core A.7 interior A.8 core A.9 rim A.10 rim A.11 core A.12 rim
Th (ppm)
62 59 79 202 53 74 73 228 85 168 55 82 518 43 59 56 91 164 54 65 426 58 31 52 41
U (ppm)
Salzwedel 2/64 sal-1 171 sal-2 191 sal-3 151 sal-1b 147 sal-4 162 sal-4b 639 sal-5 195 sal-6 234 sal-6b 21 sal-7 311 sal-8a 168 sal-8b 366 sal-9 1214 sal-10a 184 sal-10b 102 sal-11 64 sal-12a 231 sal-12b 385 sal-13a 136 sal-13b 227 sal-14 492 sal-15a 166 sal-15b 54 sal-16a 201 sal-16b 136
Drilling spot
63 24 17 113 27 39 18 29 32 125 40 52
8 9 7 28 8 102 9 69 11 15 8 46 66 8 24 18 11 19 53 12 26 8 15 27 7
Pb (ppm)
0.058 0 0.07 4.7 0 0.16 0.093 0 0 1.6 0.083 0.041
0.00 0.07 0.13 4.72 0.86 0.01 0.32 0.22 0.00 0.01 0.00 0.18 0.14 0.19 0.00 0.47 0.06 0.00 0.04 0.15 0.20 0.06 0.12 0.49 0.35
f (%)*
238
0.0937 0.047 0.0521 0.2194 0.0479 0.0518 0.0537 0.5177 0.0477 0.0501 0.0719 0.0461
0.0468 0.0462 0.0462 0.1313 0.0467 0.1676 0.0468 0.2463 0.2566 0.0454 0.0465 0.1275 0.0528 0.0467 0.2177 0.2441 0.0469 0.0477 0.3642 0.0545 0.0466 0.0469 0.2517 0.1347 0.0489
0.3456 0.3348 0.3295 0.8089 0.3445 1.6606 0.3138 3.0795 3.0373 0.3296 0.3478 1.1775 0.3808 0.3313 2.3843 2.9780 0.3375 0.3407 6.5515 0.3936 0.3230 0.3299 3.3612 1.3639 0.3408
235
Pb/ U
207
74 0.7593 39 0.3483 47 0.3749 179 2.1593 40 0.351 42 0.3673 47 0.386 626 12.5475 39 0.3439 38 0.3623 58 0.5548 36 0.3348
5 5 5 17 5 17 5 26 61 5 5 14 4 5 25 41 5 5 41 5 4 5 40 17 6
±
985 513 941 3064 507 557 707 19466 467 412 802 616
81 101 137 831 120 228 110 716 1193 142 82 238 68 105 386 1233 109 61 987 108 110 121 1308 331 135
±
0.0588 0.0538 0.0522 0.0714 0.0531 0.0514 0.0521 0.1758 0.0523 0.0525 0.056 0.0526
0.0535 0.0526 0.0518 0.0447 0.0535 0.0719 0.0487 0.0907 0.0858 0.0527 0.0542 0.0670 0.0523 0.0514 0.0794 0.0885 0.0521 0.0518 0.1305 0.0523 0.0503 0.0510 0.0968 0.0734 0.0506
206
Pb/ Pb
207
55 60 117 76 57 61 78 146 52 40 61 82
9 13 19 44 16 5 15 17 24 20 10 10 7 14 7 31 15 6 11 12 15 17 32 14 17
±
238
577 296 327 1278 302 326 337 2689 301 315 447 291
295 291 291 795 294 999 295 1419 1472 286 293 774 332 295 1270 1408 296 300 2002 342 294 295 1447 815 308
206
Pb/ U age (Ma)
4 2 3 9 2 3 3 27 2 2 4 2
4 3 3 10 3 10 3 14 32 3 4 8 3 3 14 21 3 3 20 4 3 4 21 10 4
± (Ma)
TABLE A1. COMPLETE DATA SET OF SHRIMP ANALYSES (continued) Pb/ U
206
574 303 323 1168 305 318 331 2646 300 314 448 293
301 293 289 602 301 994 277 1428 1417 289 303 790 328 291 1238 1402 295 298 2053 337 284 289 1495 874 298
207
235
Pb/ U age (Ma)
6 4 7 10 4 4 5 15 4 3 5 5
6 8 11 47 9 9 9 18 30 11 6 11 5 8 12 31 8 5 13 8 9 9 30 14 10
± (Ma)
558 363 295 969 334 259 290 2614 297 306 452 314
351 312 275 0 352 982 133 1440 1335 314 380 837 298 259 1183 1393 292 277 2104 300 208 242 1564 1026 222
206
Pb/ Pb age (Ma)
207
20 25 51 22 24 27 34 14 23 17 24 36
42 60 88 54 71 16 73 37 55 91 42 32 33 65 19 69 67 31 15 55 73 77 63 39 82
± (Ma)
103 82 111 132 90 126 116 103 101 103 99 93 Continued
84 93 106 0 84 102 222 99 110 91 77 92 111 114 107 101 101 108 95 114 141 122 93 79 139
Concordance (%)
186 Breitkreuz et al.
U (ppm)
Tuchen 1/74 (continued) A-2-1 core 80 A-2-2 core 161 A-2-3 core 109 A-2-4 core 195 A-2-5 core 37 A-2-6 core 548 A-2-7 core 187 A-2-8 core 277 A-2-9 core 67 A-2-10 core 197 A-2-11 core 132 A-2-12 rim 497 A-2-13 core 40 A-2-14 rim 229 A-2-15 core 128 A-2-16 interior 966 A-2-17 interior 157 A-2-18 rim 559 A-2-19 rim 408 A-2-20 core 79 A-2-21 interior 384 A-2-22 core 60 A-2-23 core 48 A-2-24 core 774 A-2-25 core 358 A-2-26 core 38 A-2-27 core 161 A-2-28 core 536 A-2-29 core 334 A-2-30 core 142 A-2-31 core 174 A-2-32 core 46 A-2-33 core 151 A-2-34 core 222 A-2-36 core 531 A-2-36b core 232 A-2-37 rim 415 A-2-38 core 113 A-2-39 core 128
Drilling spot
44 31 35 64 52 296 44 113 62 163 43 32 28 40 40 116 168 102 131 27 164 21 29 68 90 51 54 105 217 76 109 14 57 30 387 76 118 92 67
Th (ppm) 22 30 6 29 12 137 10 37 25 64 12 21 11 12 6 224 48 25 19 30 64 12 9 34 20 7 55 27 84 7 9 8 13 27 35 11 20 12 9
Pb (ppm) 0.71 0 0.017 0 0 0.62 0.23 1 0.079 0.078 0.13 0 0 5.2 0.069 0.21 0.5 0.013 0.11 0.47 0.042 0.68 0.26 0.016 0 0.3 0.15 0.001 1.6 0.74 0.2 0.18 0.25 0.39 0 0.13 0 0.17 0.53
f (%)*
238
0.248 0.1883 0.0514 0.1497 0.2411 0.2293 0.0526 0.126 0.3129 0.2817 0.0927 0.0471 0.2497 0.0451 0.0514 0.2354 0.2502 0.0472 0.0477 0.3552 0.1621 0.1886 0.1795 0.0472 0.0574 0.1482 0.3241 0.0521 0.2175 0.0478 0.0455 0.1628 0.0817 0.1267 0.0594 0.0461 0.0484 0.097 0.0654
317 198 68 127 373 170 59 110 392 220 101 37 365 56 72 148 229 40 39 419 122 250 275 33 52 272 336 41 154 63 55 258 107 133 46 50 46 118 85
±
3.1415 2.2969 0.3987 1.4726 2.9756 2.6052 0.3637 1.1288 4.8297 3.8717 0.7783 0.3365 3.1483 0.3296 0.352 3.3088 3.0537 0.3357 0.3316 6.9292 1.6014 2.0043 2.1458 0.3518 0.4343 1.4058 5.1182 0.374 2.4497 0.3458 0.3328 1.493 0.7695 1.1077 0.4386 0.3326 0.3589 0.8628 0.48
235
Pb/ U
207
8606 4825 1812 3019 8320 3732 1349 2479 14347 6389 2467 665 8258 1546 1757 3061 7713 785 917 13272 2611 6586 8438 608 943 11850 9369 819 4039 2013 1780 6018 2549 2523 795 1291 894 4021 2449
±
0.0919 0.0885 0.0563 0.0713 0.0895 0.0824 0.0502 0.065 0.1119 0.0997 0.0609 0.0519 0.0914 0.053 0.0497 0.1019 0.0885 0.0516 0.0504 0.1415 0.0716 0.0771 0.0867 0.0541 0.0549 0.0688 0.1145 0.0521 0.0817 0.0524 0.0531 0.0665 0.0683 0.0634 0.0536 0.0523 0.0538 0.0645 0.0532
206
Pb/ Pb
207
208 150 235 126 192 94 170 124 284 135 173 89 183 231 229 62 199 107 128 194 97 219 298 81 102 550 159 102 115 289 268 234 197 120 83 188 118 281 254
±
238
1428 1112 323 899 1392 1331 330 765 1755 1600 571 296 1437 284 323 1363 1439 297 300 1960 969 1114 1064 297 360 891 1810 327 1269 301 287 972 506 769 372 291 304 597 408
206
Pb/ U age (Ma) 16 11 4 7 19 9 4 6 19 11 6 2 19 3 4 8 12 2 2 20 7 14 15 2 3 15 16 3 8 4 3 14 6 8 3 3 3 7 5
± (Ma)
TABLE A1. COMPLETE DATA SET OF SHRIMP ANALYSES (continued) Pb/ U
206
1443 1211 341 919 1401 1302 315 767 1790 1608 584 295 1445 289 306 1483 1421 294 291 2102 971 1117 1164 306 366 891 1839 323 1257 302 292 928 579 757 369 292 311 632 398
207
235
Pb/ U age (Ma) 21 15 13 12 21 11 10 12 25 13 14 5 20 12 13 7 19 6 7 17 10 22 27 5 7 50 16 6 12 15 14 25 15 12 6 10 7 22 17
± (Ma) 1465 1393 463 967 1415 1255 202 773 1831 1618 636 279 1456 329 180 1660 1394 268 214 2245 976 1123 1354 375 407 892 1873 290 1238 305 332 823 878 722 353 297 364 758 339
206
Pb/ Pb age (Ma)
207
43 32 93 36 41 22 79 40 46 25 61 39 38 99 107 11 43 48 59 24 28 57 66 34 42 166 25 45 28 126 115 73 60 40 35 82 49 92 108
± (Ma)
98 80 70 93 98 106 163 99 96 99 90 106 99 87 180 82 103 111 140 87 99 99 79 79 88 100 97 113 102 99 86 118 58 106 105 98 84 79 120 Continued
Concordance (%)
Chronostratigraphic structure of Far Eastern Avalonia 187
113 76 120 169 80 58 93 171 73 50 87 78 158 66 74 199 224 187 162 200 209
Wysoka Kamienska-2 wys-1 215 wys-2a 191 wys-2b 228 wys-3 311 wys-4 201 wys-4b 170 wys-5 211 wys-6 326 wys-7 185 wys-8 142 wys-9 216 wys-10 183 wys-11 281 wys-12 179 wys-13 185 wys-14 362 wys-15 407 wys-16 321 wys-17 452 wys-18 343 wys-19 339 11 9 11 15 11 8 11 16 10 7 11 9 14 9 9 18 20 16 23 18 18
Pb (ppm) 0.20 0.18 0.14 0.10 4.14 0.03 2.00 0.03 4.65 0.00 1.79 0.02 0.24 4.52 0.16 0.44 0.55 0.00 2.82 0.30 0.74
f (%)*
Zdrój-1 B.1 interior 309 103 15 1.2 B.2 core 70 27 3 0.22 B.3 core 216 140 11 0.1 B.4 interior 83 47 4 0.62 B.5 rim 93 42 5 0.23 B.6 interior 90 48 5 4.6 B.7 core 215 136 70 0.29 B.8 core 198 114 12 2 B.9 core 466 329 24 0 B.10 core 269 97 53 0.06 B.11 interior 548 23 45 0.69 B.12 interior 447 224 125 0.15 B.13 interior 529 399 28 0.95 B.14 interior 87 36 4 0.22 B.15 core 232 131 12 0 B.16 core 106 62 5 0.4 B.17 rim 133 67 7 1.8 B.18 interior 101 47 5 0.61 B.19 interior 98 56 5 0 Note: Data are summarized in Table 1. 206 *Percentage of Pb that is common lead.
Th (ppm)
U (ppm)
Drilling spot
238
0.047 0.048 0.0466 0.0469 0.0471 0.0475 0.2934 0.0539 0.0471 0.1926 0.0861 0.2609 0.047 0.0458 0.0474 0.0474 0.0466 0.0457 0.0471
0.0469 0.0459 0.0465 0.0463 0.0474 0.0467 0.0472 0.0474 0.0468 0.0464 0.0473 0.0479 0.0479 0.0454 0.0472 0.0467 0.0465 0.0471 0.0466 0.0477 0.0473
44 67 47 61 58 60 272 54 52 158 69 218 48 59 56 59 56 80 68
4 4 4 5 4 4 4 4 4 5 9 10 9 5 5 4 4 5 4 5 5
±
0.3344 0.3394 0.3412 0.3083 0.3472 0.3269 3.8762 0.4196 0.348 2.0482 0.7117 3.353 0.3176 0.33 0.351 0.3199 0.3334 0.3365 0.3281
0.3269 0.3275 0.3322 0.3296 0.3410 0.3355 0.3389 0.3390 0.3362 0.3370 0.3284 0.3456 0.3353 0.3214 0.3295 0.3150 0.3386 0.3498 0.3212 0.3320 0.3392
235
Pb/ U
207
893 1710 1178 1710 1491 2479 5961 1409 666 2725 929 3997 1153 1502 788 1565 1477 2157 1392
92 88 90 98 118 90 100 79 116 68 125 125 117 152 120 102 97 66 100 113 113
±
0.0516 0.0513 0.0531 0.0477 0.0535 0.05 0.0958 0.0565 0.0536 0.0771 0.06 0.0932 0.049 0.0522 0.0537 0.049 0.0519 0.0535 0.0505
0.0506 0.0518 0.0518 0.0517 0.0522 0.0521 0.0521 0.0519 0.0521 0.0526 0.0504 0.0523 0.0507 0.0513 0.0506 0.0489 0.0528 0.0539 0.0500 0.0505 0.0520
206
Pb/ Pb
207
123 239 169 250 212 366 107 174 77 73 56 70 165 220 95 224 214 318 193
12 12 12 13 16 12 13 10 16 8 14 14 13 22 16 14 13 7 14 15 15
±
295 289 293 292 299 294 297 298 295 293 298 302 302 286 297 294 293 297 294 300 298
238
296 302 294 295 296 299 1659 338 297 1136 532 1495 296 289 299 298 294 288 297
206
Pb/ U age (Ma)
3 4 3 4 4 4 14 3 3 9 4 11 3 4 3 4 3 5 4
3 3 3 3 3 3 3 3 3 3 6 6 6 3 3 3 3 3 3 3 3
± (Ma)
TABLE A1. COMPLETE DATA SET OF SHRIMP ANALYSES (continued) Pb/ U
206
293 297 298 273 303 287 1609 356 303 1132 546 1493 280 290 305 282 292 295 288
287 288 291 289 298 294 296 296 294 295 288 301 294 283 289 278 296 305 283 291 297
207
235
Pb/ U age (Ma)
7 13 9 13 11 19 12 10 5 9 6 9 9 11 6 12 11 16 11
7 7 7 8 9 7 8 6 9 5 10 9 9 12 9 8 7 5 8 9 9
± (Ma)
266 254 334 85 351 193 1544 471 355 1124 602 1492 148 296 358 148 281 348 219
222 275 277 270 293 291 289 282 289 314 211 298 229 255 223 142 322 366 195 218 286
206
Pb/ Pb age (Ma)
207
55 107 72 118 90 163 21 68 32 19 20 14 77 96 40 104 94 135 88
58 54 56 60 73 54 61 46 72 36 69 62 60 102 77 68 59 32 66 71 70
± (Ma)
111 119 88 346 84 155 107 72 84 101 88 100 200 98 83 202 105 83 136
133 105 106 108 102 101 103 106 102 93 141 101 132 112 133 207 91 81 151 138 104
Concordance (%)
188 Breitkreuz et al.
Chronostratigraphic structure of Far Eastern Avalonia REFERENCES CITED Abragawa, M.A., 1993, Petrology and geochemistry of the Rotliegendes volcanic rocks in Denmark and their tectonic implications, in Stouge, S., and Vejbæk, O.V., eds., EFP-89; Område 1: Olie og Naturgas: Dynamisk/ stratigrafisk analyse af Palæozoikum i Danmark: Copenhagen, Danmarks Geologiske Undersøgelse (DGU) v. 3, Kunderapport 35, 335 p. André, L., Hertogen, J., and Deutsch, S., 1986, Ordovician-Silurian magmatic provinces in Belgium und the Caledonian orogeny in middle Europe: Geology, v. 14, p. 879–882, doi: 10.1130/0091-7613(1986)14<879: OMPIBA>2.0.CO;2. Arthaud, F., and Matte, P., 1977, Late Paleozoic strike-slip faulting in southern Europe and northern Africa: Results of right-lateral shear zone between the Appalachians and the Urals: Geological Society of America Bulletin, v. 88, p. 1305–1320, doi: 10.1130/0016-7606(1977)88<1305: LPSFIS>2.0.CO;2. Awdankiewicz, M., Breitkreuz, C., and Ehling, B.-C., 2004, Emplacement textures in Late Palaeozoic andesite sills of the Flechtingen-Roßlau block, north of Magdeburg (Germany), in Breitkreuz, C., and Petford, N., eds., Physical geology of high-level magmatic systems: London, Geological Society of London Special Publication 234, p. 51–66. Banka, D., Pharaoh, T.C., Williamson, J.P., and TESZ Project Potential Field Core Group, 2002, Potential field imaging of Palaeozoic orogenic structure in northern and central Europe: Tectonophysics, v. 360, p. 23–45. Baumann, A., Grauert, B., Mecklenburg, S., and Vinx, R., 1991, Isotopic age determinations of crystalline rocks of the Upper Harz Mountains, Germany: Geologische Rundschau, v. 80, p. 669–690, doi: 10.1007/ BF01803694. Bayer, U., Grad, M., Pharaoh, T.C., Thybo, H., Guterch, A., Banka, D., Lamarche, J., Lassen, A., Lewerenz, B., Scheck, M., and Marotta, A.-M., 2002, The southern margin of the East European craton: New results from seismic sounding and potential fields between the North Sea and Poland: Tectonophysics, v. 360, p. 301–314, doi: 10.1016/S0040-1951(02)00359-1. Benek, R., Paech, H.J., and Schirmer, B., 1973, Zur Gliederung der permosilesischen Vulkanite der Flechtinger Scholle: Zeitschrift der geologischen Wissenschaften, v. 1, p. 867–878. Benek, R., Kramer, W., McCann, T., Scheck, M., Negendank, J.F.W., Korich, D., Huebscher, H.-D., and Bayer, U., 1996, Permo-Carboniferous magmatism of the northeast German basin: Tectonophysics, v. 266, p. 379–404, doi: 10.1016/S0040-1951(96)00199-0. Brecht, G., 1999, Authigene Phyllosilikate in permokarbonen SiO2-reichen Vulkaniten Ostdeutschlands: Berliner geowissenschaftliche Abhandlungen, v. A201, p. 1–181. Breitkreuz, C., and Kennedy, A., 1999, Magmatic flare-up at the Carboniferous/Permian boundary in the NE German basin revealed by SHRIMP zircon ages: Tectonophysics, v. 302, p. 307–326, doi: 10.1016/S00401951(98)00293-5. Breitkreuz, C., Karnkowski, P., Muszynski, A., Panczyk, M., and Protas, A., 2000, The Wolsztyn Ridge in western Poland during the Early Permian: Paleogeography, volcanic activity and coarse-grained sedimentation at an intra-basinal high: Exkursionsführer und Veröffentlichungen der Gesellschaft für Geologische Wissenschaften, v. 209, p. 58–60. Franke, W., 2000, The mid-European segment of the Variscides: Tectonostratigraphic units, terrane boundaries and plate tectonic evolution, in Franke, W., Haak, V., Oncken, O., and Tanner, D., eds., Orogenic processes— Quantification and modelling in the Variscan belt: London, Geological Society of London Special Publication 179, p. 35–61. Gáal, G., and Gorbatschev, R., 1987, Precambrian geology and evolution of the Central Baltic Shield: Precambrian Research, v. 35, p. 1–382, doi: 10.1016/0301-9268(87)90043-X. Gaitzsch, B., 1995, Extramontane Senken im variscischen Finalstadium in Norddeutschland—Lithofaziesmuster, Tektonik und Beckenentwicklung [Ph.D. thesis]: Freiberg, Germany, Technische Universität Freiberg, 101 p. Geißler, M., and Breitkreuz, C., 2004, Vulkanische Faziestypen und ihre zeitlich-räumliche Verteilung im Permokarbon des Zentraleuropäischen Bekkensystems: Schriftreihe der Deutschen Gesellschaft für Geowissenschaften, v. 34, p. 42. Heeremans, M., and Faleide, J.I., 2004, Late Carboniferous–Permian tectonics and magmatic activity in the Skagerrak, Kattegat and the North Sea, in Wilson, M., ed., Permo-Carboniferous magmatism and rifting in Europe: London, Geological Society of London Special Publication 223, p. 157–176.
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Heeremans, M., Faleide, J.I., and Larsen, B.T., 2004, Late Carboniferous-Permian of NW Europe; an introduction to a new regional map, in Wilson, M., ed., Permo-Carboniferous magmatism and rifting in Europe: London, Geological Society of London Special Publication 223, p. 75–88. Hoffmann, N., and Stiewe, H., 1994, Neuerkenntnisse zur geologisch-geophysikalischen Modellierung der Pritzwalker Anomalie im Bereich des Ostelbischen Massivs: Zeitschrift für geologische Wissenschaften, v. 22, p. 161–171. Huckriede, H., Wemmer, K., and Ahrendt, H., 2004, Palaeogeography and tectonic structure of allochthonous units in the German part of the RhenoHercynian belt (central European Variscides): International Journal of Earth Sciences, v. 93, p. 414–431, doi: 10.1007/s00531-004-0397-4. Jackowicz, E., 1994, Persmkie skaly wulkaniczne pólnocnej czesci monokliny przedsudeckiej: Prace Pan´stwowego Instytutu Geologicznego Warszawa, v. 145, p. 1–47. Korich, D., 1992, Zur Vulkanologie und Korrelation der permosilesischen Vulkanite im Darß-Uckermark-Eruptivkomplex/Nordostdeutschland: Zeitschrift für geologische Wissenschaften, v. 20, p. 467–473. Linnemann, U., and Romer, R.L., 2002, The Cadomian orogeny in SaxoThuringia, Germany: Geochemical and Nd-Sr-Pb isotopic characterisation of marginal basins with constraints to geotectonic setting and provenance: Tectonophysics, v. 352, p. 33–64, doi: 10.1016/S0040-1951(02)00188-9. Linnemann, U., Romer, R.L., Gehmlich, M., and Drost, K., 2003, Paläogeographie und Provenance des Saxothuringikums unter besonderer Berücksichtigung der Geochronologie von prävariszischen Zirkonen und der Nd-Isotopie von Sedimenten: Geologica Saxonica, v. 48/49, p. 121–132. Lippolt, H.J., Raczek, I., and Schleicher, H., 1982, Isotopenalter (40Ar/39Ar; RbSr) eines Unteren Rotliegend-Biotits aus der Bohrung Wrzesnia/Polen: Der Aufschluß, v. 33, p. 13–25. Lorenz, V., and Nicholls, I.A., 1984, Plate and intraplate processes of Hercynian Europe during the Late Paleozoic: Tectonophysics, v. 107, p. 25–56, doi: 10.1016/0040-1951(84)90027-1. Maliszewska, A., Kiersnowski, H., and Jackowicz, E., 2003, Wulkanoklastyczne osady czerwonego spagowca dolnego na obszarze wielkopolski: Prace Pan´stwowego Instytutu Geologicznego Warszawa, v. 179, p. 1–59. Marx, J., Huebscher, H.-D., Hoth, K., Korich, D., and Kramer, W., 1995, Vulkanostratigraphie und Geochemie der Eruptivkomplexe, Stratigraphie von Deutschland I—Norddeutsches Rotliegendbecken: Courier Forschungshefte Senckenberg Frankfurt, v. 183, p. 54–83. Matte, P., 2001, The Variscan collage and orogeny (480–290 Ma) and the tectonic definition of the Armorica mircroplate: A review: Terra Nova, v. 13, p. 122–128, doi: 10.1046/j.1365-3121.2001.00327.x. McCann, T., 1998, Lower Paleozoic evolution of the northeast German Basin/ Baltica borderland: Geological Magazine, v. 135, p. 129–142, doi: 10.1017/S0016756897007863. McNamara, A.K., Niocaill, C.M., van den Pluijm, B.A., and van den Voo, R., 2001, West African proximity of the Avalon terrane in the latest Precambrian: Geological Society of America Bulletin, v. 113, p. 1161–1170. Meissner, R., and Krawczyk, C.H., 1999, Caledonian and Proterozoic terrane accretion in the southwest Baltic Sea: Tectonophysics, v. 314, p. 255–267, doi: 10.1016/S0040-1951(99)00247-4. Murphy, J.B., Keppie, J.D., Dostal, J., and Nance, R.D., 1999, Neoproterozoic– Early Paleozoic evolution of Avalonia, in Ramos, V., ed., Laurentia-Gondwana connections before Pangea: Boulder, Colorado, Geological Society of America Special Paper 336, p. 253–266. Murphy, J.B., Nance, R.D., and Keppie, J.D., 2002, West Africa proximity of the Avalon terrane in latest Precambrian: Discussion: Geological Society of America Bulletin, v. 114, p. 1049–1050, doi: 10.1130/00167606(2002)114<1049:DARWAP>2.0.CO;2. Nance, R.D., and Murphy, J.B., 1994, Contrasting basement isotopic signatures and the palinspastic restoration of peripheral orogens: Example from the Neoproterozoic Avalonian-Cadomian belt: Geology, v. 22, p. 617–620, doi: 10.1130/0091-7613(1994)022<0617:CBISAT>2.3.CO;2. Nesbor, H.D., Buggisch, W., Flick, H., Horn, M., and Lippert, H.-J., 1993, Vulkanismus im Devon des Rhenoherzynikums: Fazielle und paläoeographische Entwicklung vulkanisch geprägter mariner Becken am Beispiel des Lahn-Dill-Gebietes: Geologische Abhandlungen Hessen, v. 98, p. 3–87. Neumann, E.-R., Wilson, M., Heeremans, M., Spencer, E.A., Obst, K., Timmerman, M.J., and Kirstein, L., 2004, Late Carboniferous–Permian rifting and magmatism in southern Scandinavia, the North Sea and northern Germany: A review: London, Geological Society of London Special Publication 223, p. 11–40.
190
Breitkreuz et al.
Noble, S.R., Tucker, R.D., and Pharao, T.C., 1993, Lower Palaeozoic and Precambrian igneous rocks from eastern England, and their bearing on late Ordovician closure of the Tornquist Sea: Constraints from U-PB and Nd isotopes: Geological Magazine, v. 130, p. 835–846. Obst, K., Hammer, J., Katzung, G., and Korich, D., 2004, The Mesoproterozoic basement in the southern Baltic Sea: Insights from the G 14-1 off-shore borehole: International Journal of Earth Sciences, v. 93, p. 1–12, doi: 10.1007/s00531-003-0371-6. Paulick, H., and Breitkreuz, C., 2005, The Late Paleozoic felsic lava–dominated large igneous province in north east Germany: Volcanic facies analysis based on drill cores: International Journal of Earth Sciences, v. 94, p. 834–850, doi: 10.1007/s00531-005-0017-y. Pharaoh, T.C., 1999, Paleozoic terranes and their lithospheric boundaries within the Trans-European suture zone (TESZ): A review: Tectonophysics, v. 314, p. 17–41, doi: 10.1016/S0040-1951(99)00235-8. Romer, R.L., Förster, H.-J., and Breitkreuz, C., 2001, Intracontinental extensional magmatism with a subduction fingerprint: The late Carboniferous Halle volcanic complex (Germany): Contributions to Mineralogy and Petrology, v. 140, p. 201–221. Scheck-Wenderoth, M., and Lamarche, J., 2005, Crustal memory and basin evolution in the Central European Basin System—New insights from a 3D structural model: Tectonophysics, v. 397, p. 143–165, doi: 10.1016/ j.tecto.2004.10.007. Smith, R.B., and Braile, L.W., 1994, The Yellowstone hotspot: Journal of Volcanology and Geothermal Research, v. 61, p. 121–187, doi: 10.1016/03770273(94)90002-7.
Stemmerik, L., Ineson, J.R., and Mitchell, J.G., 2000, Stratigraphy of the Rotliegend Group in the Danish part of the northern Permian basin, North Sea: Journal of the Geological Society of London, v. 157, p. 1127–1136. Tucker, R.D., and Pharaoh, T.C., 1991, U-Pb zircon ages for Late Precambrian igneous rocks in southern Britain: Journal of the Geological Society of London, v. 148, p. 435–443. Williamson, J.P., Pharaoh, T.C., Banka, D., Thybo, H., Laigle, M., and Lee, M.K., 2002, Potential field modelling of the Baltica–Avalonia (Thor– Tornquist) suture beneath the southern North Sea: Tectonophysics, v. 360, p. 47–60, doi: 10.1016/S0040-1951(02)00346-3. Winchester, J.A., Pharaoh, T.C., and Verniers, J., 2002, Paleozoic amalgamation of central Europe: An introduction and synthesis of new results from recent geological and geophysical investigations, in Winchester, J.A., ed., Palaeozoic amalgamation of central Europe: London, Geological Society of London Special Publication 201, p. 1–18. Ziegler, P., 1990, Geological atlas of western and central Europe, 2nd edition: The Hague, Shell Internationale Petroleum Maatschappij, 239 p. Zwingmann, H., Clauer, N., and Gaupp, R., 1998, Timing of fluid flows in a sandstone reservoir of the northern German Rotliegend (Permian) by KAr dating of related hydrothermal illite, in Parnell, J., ed., Dating and duration of fluid flow and fluid-rock interaction: London, Geological Society of London Special Publication 144, p. 91–106.
MANUSCRIPT ACCEPTED BY THE SOCIETY 3 OCTOBER 2006
Printed in the USA
Geological Society of America Special Paper 423 2007
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) Kerstin Drost* Staatliche Naturhistorische Sammlungen Dresden, Museum für Mineralogie und Geologie, Königsbrücker Landstrasse 159, 01109 Dresden, Germany Rolf L. Romer GeoForschungsZentrum Potsdam, Telegrafenberg, 14473 Potsdam, Germany Ulf Linnemann Staatliche Naturhistorische Sammlungen Dresden, Museum für Mineralogie und Geologie, Königsbrücker Landstrasse 159, 01109 Dresden, Germany Oldrˇich Fatka Petr Kraft Jaroslav Marek Charles University Prague, Institute of Geology and Paleontology, Albertov 6, 128 43 Prague 2, Czech Republic
ABSTRACT Nd-Sr-Pb isotope data are used to characterize the sources of Late Neoproterozoic and Early Paleozoic siliciclastic rocks of the Teplá-Barrandian unit of the Bohemian Massif. Geochemical and isotopic signatures of samples from different stratigraphic levels reflect changing sources and weathering conditions through time and allow a correlation with shifting geotectonic regimes. Late Neoproterozoic rocks were deposited in a magmatic arc–related setting within the Avalonian-Cadomian belt at the periphery of West Gondwana. Fine-grained graywackes yield crustal residence ages (TDM) of 2.17–1.49 Ga, documenting contributions of old crust. Their εNd570 values, as well as Pb and Sr isotopic compositions, reflect mixing of detritus derived from old crust with a Neoproterozoic magmatic arc component. The change in the geotectonic regime to transtension/rifting occurred during the terminal Neoproterozoic and is documented by more radiogenic εNdT values (−6.0 to +1.0) and younger TDM *E-mail:
[email protected]. Drost, K., Romer, R.L., Linnemann, U., Fatka, O., Kraft, P., and Marek, J., 2007, 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), 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. 191–208, doi: 10.1130/2007.2423(08). For permission to copy, contact
[email protected]. ©2007 Geological Society of America. All rights reserved.
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Drost et al. (1.65–1.12 Ga) of the Cambrian sediments. Besides the involvement of a post-Neoproterozoic juvenile source, the Lower Cambrian basin was also fed from an old upper crustal domain, as indicated by their high 207Pb/206Pb values. In contrast, Middle Cambrian siliciclastic rocks are mainly derived from the Cadomian basement. In the Ordovician pelites, εNdT values of −9.6 to −8.3 and radiogenic Sr and Pb isotopic compositions reflect an increasing input of material derived from the cratonic hinterland. Their TDM values range from 2.02 to 1.88 Ga. The uniform geochemical and isotopic compositions of the Ordovician samples indicate efficient mixing of the detritus prior to deposition in a mature rift or shelf environment at the Gondwanan margin. Keywords: isotope geochemistry, sedimentary rocks, Neoproterozoic, Early Paleozoic, provenance
INTRODUCTION The Teplá-Barrandian unit of the Bohemian Massif was part of the Late Neoproterozoic Avalonian-Cadomian tectonostratigraphic belt at the periphery of West Gondwana (Murphy and Nance, 1989; Nance et al., 1991; Nance and Murphy, 1994, 1996) and belongs to the Cadomian-type terranes that formed along the west African margin of the supercontinent (Murphy et al., 2004). Following the dispersal of the Avalonian-Cadomian orogen by Late Neoproterozoic to Early Paleozoic strike-slip and rifting processes, the Teplá-Barrandian unit became incorporated into the central European Variscides during the collision of Gondwana with Laurussia. Because the Teplá-Barrandian unit was situated in an upper crustal level throughout the entire Paleozoic, it was largely spared from Variscan metamorphism. The boundaries of the Teplá-Barrandian unit with the adjacent Saxo-Thuringian and Moldanubian high-grade units are represented by major fault and shear zones (Zulauf et al., 1997a, 2002, and references therein). The metamorphic grade within the Teplá-Barrandian unit increases toward the west and northwest (e.g., Vejnar, 1966) and is predominantly related to Cadomian orogenic processes (Pašava and Amov, 1993; Dörr et al., 1998; Zulauf et al., 1999, and references therein). The eastern and southeastern parts of the Teplá-Barrandian unit are composed of very-low-grade to low-grade Neoproterozoic and unmetamorphosed Cambrian to Middle Devonian overstep sequences (Fig. 1). Neoproterozoic to Paleozoic siliciclastic sedimentary rocks of peri-Gondwanan crustal fragments were subject to various studies using Nd and Sr isotopic data for provenance analyses (e.g., Beetsma, 1995; Nägler et al., 1995; Simien et al., 1999; Linnemann and Romer, 2002; Murphy and Nance, 2002). The composition of radiogenic isotopes (e.g., Sr, Nd, Hf, Pb) depends on the time-integrated parent-to-daughter ratio (e.g., 87Rb/86Sr, 147Sm/144Nd, 176Lu/177Hf), which differs between mantle and crust and among different crustal reservoirs, and therefore gives information on the provenance of terrigenous siliciclastic rocks and variations of their sources through time. The composition of radiogenic isotopes, therefore, may contribute to unraveling potential rift-drift-docking histories of
terranes (e.g., Thorogood, 1990; McCaffrey, 1994; Linnemann et al., 2004). In this study we investigated the geochemical and Nd-Sr-Pb isotopic signatures of the Late Neoproterozoic Cadomian basement and the Early Paleozoic siliciclastic overstep sequences of the Teplá-Barrandian unit in order to detect changes in provenance during this period. We demonstrate the development of the Teplá-Barrandian unit from a Cadomian magmatic arc–related setting to Early Paleozoic rift and shelf environments. As different reservoirs, such as the depleted mantle, the lower continental crust, and the upper continental crust, are characterized by specific geochemical and isotopic compositions, the signatures of siliciclastic sediments allow the identification of varying contributions of material derived from these contrasting sources. Because detrital contributions from diverse sources produce mixed isotopic compositions in the resulting sediments, the combination of differently behaving isotopic systems allows more detailed conclusions on contributing sources. GEOLOGICAL FRAMEWORK Neoproterozoic The lithostratigraphic subdivision of the Teplá-Barrandian Neoproterozoic rocks is chiefly based on the presence (or absence) and the character of synsedimentary volcanic rocks (e.g., Kettner, 1918; Röhlich, 1965; Cháb, 1993). The division into an older Kralupy-Zbraslav Group and a younger Šteˇ chovice Group (e.g., Mašek, 2000; Fig. 2) was established principally based on studies of Neoproterozoic successions east of the Barrandian Paleozoic (Fig. 1), whereas the large area southwest, west, and northwest of the Paleozoic rocks is completely assigned to the Blovice Formation of the KralupyZbraslav Group (Cháb, 1993; Chaloupský et al., 1995; Cháb et al., 1997). Microfossils confirm an Upper Riphean to Vendian age (corresponding to the Ediacaran in the current terminology of Gradstein et al., 2005) for both lithostratigraphic units and allow a correlation with the Brioverian of the Armorican Massif in northwestern France (Konzalová, 1981; Pacltová, 1990; Fatka and Gabriel, 1991). The Kralupy-Zbraslav Group is composed
Nd-Sr-Pb isotopic signatures of Neoproterozoic–Early Paleozoic siliciclastic rocks
A A
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B B
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Praha
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D
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DB1/1 DB1/2 STA DB1/5 DB6/1 Skryje DB6/3
KR VC
50°00'
Rˇ ícˇany
ViBe1A KDLeS2 Beroun KoLeS1Konìprusy Karlštejn KoLeS2 Srbsko Zdice
PLZENˇ
Rokycany
NPZb4 NPZb5 Jílové u Prahy
Variscan granitoids Metamorphic Neoproterozoic + Paleozoic of the Islet zone
Mníšek pod Brdy
Horˇovice Hoøovice
J4 J8
Mýto
KlRS1
LeZb4
Rˇevnice Øevnice Suchomasty
Žebrák ?ebrák Radnice
Jince
SVC
MM2
PJB
14° 40' Úvaly
DB S1 DB D2
Neogene
Dobrˇíš
Upper Cretaceous
Dobrˇíš Dobrˇíš 02b
Upper Carboniferous and Permian Devonian marine sediments
Prˇíbram Pøíbram Sedlcˇany Sedlèany
Silurian marine sediments and volcanics Ordovician marine sediments and volcanics Upper Cambrian volcanics Cambrian continental and marine sediments
Blovice
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0
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Neoproterozoic marine sediments and volcanics
Figure 1. (A) Distribution of Variscan massifs in central Europe. (B) Location of the study area within the Bohemian Massif. (C) Simplified geological map of the Barrandian (modified after Chlupácˇ et al., 1998) with sampling localities. KRVC—Krˇivoklát-Rokycany volcanic complex; PJB—Prˇíbram-Jince basin; STA—Skryje-Týrˇovice area; SVC—Strašice volcanic complex.
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Upper Mid. Low.
Vinice F. Letná F. Dobrotivá F. Šárka F.
Klabava F.
* erosional disconformity KRVC SVC Pavlovsko
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Chumava-Baština F.
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Libenˇ F.
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Sádek F. Zˇitec-Hluboš F.
Davle F. Blovice Fm.
Kralupy-Zbraslav Group
Late Neoproterozoic
Šteˇchovice Group
Cadomian angular unconformity
glaciomarine diamictite
basic volcanics
oolithic iron ore
intermediate volcanics
shales sandstones/ quartzites
acid volcanics
conglomerates
black shales
cherts
siliciclastic alternations
pyroclastics and tuffites
Figure 2. Stratigraphy of the Barrandian Neoproterozoic and Early Paleozoic volcanosedimentary sequences. Asterisks (*) stand for the Middle Cambrian continental Ohrazenice Formation, as well as for the Lower Ordovician transgressive Trˇenice Formation that is partly overlain by cherts of the Mílina Formation. Fm, F—formation; KRVC—KrˇivoklátRokycany volcanic complex; SVC—Strašice volcanic complex. Compiled from Havlícˇek and Vaneˇk (1966), Havlícˇek (1971), Vidal et al. (1975), Chlupácˇ (1993, 1995), Kraft et al. (1999, 2004).
of alternating shales, siltstones, and sandstones with interbedded volcanics and cherts (Fig. 2). Although basic volcanics are widespread in the thick Blovice Formation in the lower part of the Kralupy-Zbraslav Group, intermediate to acidic igneous rocks and appropriate pyroclastics are typical for the overlying Davle Formation. The top of the Davle Formation is represented by silicified black shales, which pass into siltstones and shales of the Šteˇ chovice Group. The volcanic rocks of the KralupyZbraslav Group form northeast–southwest-trending belts and belong to three major series of tholeiitic, transitional, and alkaline composition, respectively. The Jílové zone in the southeast of the Teplá-Barrandian unit comprises tholeiitic volcanics in the lower part and rocks of a calc-alkaline association in the upper part (Waldhausrová, 1984). The geochemical signatures of the igneous rocks correspond to volcanic arc and back-arc geotectonic settings (Fiala, 1977, 1978; Pelc and Waldhausrová, 1994; Waldhausrová, 1984, 1997a,b). Chaloupský et al. (1995) point out that it is probable that at least parts of the Blovice and Davle formations (Fig. 2) were formed contemporaneously in an island arc–back-arc (–inter-arc) system. Alternating shales, siltstones, and graywackes are characteristic for the succeeding Šteˇ chovice Group. Volcanic activity is only indicated by thin layers of tuff and tuffite. Intercalations of conglomerates occur in the middle part (Fig. 2). Lithostratigraphic subdivisions of the Teplá-Barrandian Neoproterozoic successions differing from the one described above were proposed by several workers (for details see, e.g., Mašek, 2000; Röhlich, 2000, and references therein). According to Röhlich (2000), the Neoproterozoic basement of the TepláBarrandian unit consists of three microsegments with individual lithostratigraphic features separated by northeast–southwesttrending major faults. Holubec (1995) proposed a lithostratigraphic classification into the Rabštejn-Úslava, Zvíkovec, and Šteˇ chovice groups (top), respectively, which are separated by unconformities. The several-thousand-meters-thick Neoproterozoic siliciclastic sequences were deposited as turbidites and gravity flows on a subsiding seafloor at the northern periphery of Gondwana (Chlupácˇ, 1993). Graywacke and chert pebbles in the middle and upper parts of the Šteˇ chovice Group, as well as radiometric age data obtained from rhyolite pebbles, are interpreted to document accretion, uplift, and erosion of the Cadomian island arc. Tentative geotectonic models for the Cadomian evolution of the Teplá-Barrandian unit assume subduction accompanied by the formation of magmatic arc, back-arc basin, and remnant arc basin (Krˇ íbek et al., 2000; Dörr et al., 2002), of which the island arc activity spans at least the period from 617 ± 10 Ma to 568 ± 3 Ma (Dörr et al., 2002; Drost et al., 2004). The peak of a lowpressure/high-temperature metamorphic event with top-to-thenorth kinematics is dated ca. 550–540 Ma (Th-U-Pb chemical ages of metamorphic monazite) and interpreted to be related to arc (or similar-)continent collision and slab breakoff (Zulauf et al., 1999). The subsequent collapse of the thickened crust around the Precambrian-Cambrian boundary is expressed by normal
Nd-Sr-Pb isotopic signatures of Neoproterozoic–Early Paleozoic siliciclastic rocks faulting, causing exhumation of amphibolite-facies rocks in the western and northwestern parts of the Teplá-Barrandian unit and crustal tilting before 523 Ma (Zulauf et al., 1997b, 1999). Early Paleozoic Cambrian rocks occur in the Skryje-Týrˇovice area and in the Prˇ íbram-Jince basin (Fig. 1). Lower Cambrian sedimentary sequences crop out in the Prˇíbram-Jince basin and reach a thickness of up to 2500 m. They are lithostratigraphically subdivided into five formations (Havlícˇek, 1971). Conglomerates of the ŽitecHluboš Formation unconformably overlie Cadomian deformed basement (Fig. 2). This basal formation of the Paleozoic overstep sequence is characterized by relatively immature conglomerates and sandstones with pebbles that are interpreted to be derived from the adjacent Neoproterozoic units (Kukal, 1971). Apart from the basal Žitec-Hluboš Formation, the Lower Cambrian siliciclastic sediments are highly mature and partly mixed with material derived from synsedimentary volcanism (Kukal, 1971; Drost et al., 2004). These sediments occur in a fault-bound basin (Kukal, 1971; Havlícˇek, 1971). Lower Cambrian magmatic rocks (ca. 523–511 Ma; U-Pb zircon) include calc-alkaline plutons that were emplaced in northeast–southwest- to ENE–WSW-trending transtensive shear zones (Zulauf, 1997; Zulauf et al., 1997b; Dörr et al., 1998, and references therein) exposed in the western and northwestern parts of the Teplá-Barrandian unit. Continued subsidence resulted in a marine transgression in the Middle Cambrian (Jince Formation). In the Skryje-Týrˇovice basin, the Jince Formation unconformably overlies deformed Neoproterozoic successions. Lower Cambrian rocks are not present. In the Prˇ íbramJince basin, there is a transition from the coarser-grained, continental Chumava-Baština Formation to the fine-grained, marine Jince Formation. The sediments of the overlying Ohrazenice Formation were deposited during a regression (Havlícˇek, 1971; Kukal, 1971). The fauna of the Jince Formation shows relations to that of southwestern Europe and Morocco (Fatka et al., 1998; Alvaro et al., 2003). Persisting magmatic activity within a transtensional regime is documented by the 505-Ma-old Tis granite intruding Neoproterozoic sedimentary rocks in the northwest of the Teplá-Barrandian unit (Venera et al., 2000). The Upper Cambrian is predominantly represented by subaerial andesite-rhyolite volcanism preserved in the Strašice and Krˇ ivoklát-Rokycany volcanic complexes (SVC and KRVC in Figures 1 and 2). Geochemical studies revealed a combination of subduction-related and intraplate features (Patocˇka et al., 1993; Drost et al., 2004). Zircon crystals from rhyolites of the Krˇ ivoklát-Rokycany volcanic complex show morphologies corresponding to alkaline magmatism with high formation temperatures (after Pupin, 1980) and yielded a U-Pb sensitive highresolution ion microprobe (SHRIMP) age of 499 ± 4 Ma (Drost et al., 2004; Gehmlich and Drost, 2005). Rb-Sr isotopic studies of the same complex point to a lower crust or mantle origin of the magma (Vidal et al., 1975). The Upper Cambrian gap in sedimentation is only locally disrupted by siliciclastics of the Pavlovsko
195
Formation deposited in the Prˇ íbram-Jince basin (Havlícˇek, 1971; Kukal, 1971). The features of the Cambrian volcanosedimentary succession are compatible with a transtensional setting, commencing with the formation of a pull-apart basin. Continuing subsidence led to crustal thinning and the Middle Cambrian transgression. In the Upper Cambrian, advanced thinning of the crust gave rise to asthenospheric doming accompanied by uplift of the crust and intense subaerial volcanism (Drost et al., 2004). During the Lower Ordovician the doming of the asthenosphere was replaced by thermal subsidence accompanied by a major transgression. Ordovician siliciclastic sediments and pyroclastic rocks with intercalations of oolithic iron ores reach a thickness of ~2500 m (Havlícˇek, 1998). A “Mediterranean province” fauna indicates a cold- or cool-water environment and can be correlated with other peri-Gondwanan regions, such as Iberia, France, Sardinia, and the Italian-Austrian Carnic Alps (Havlícˇek and Vaneˇ k, 1966; Havlícˇek and Fatka, 1992; Štorch et al., 1993). Terrigenous siliciclastic sequences are developed as shallow-water sandy facies and deeper-water shales. Detrital white mica fractions from Lower to Upper Ordovician siliciclastic sedimentary rocks yielded KAr ages of 612–585 Ma and suggest a crystalline source of Late Neoproterozoic (Ediacaran) age, which was most probably represented by the deeply eroded Cadomian magmatic arc (Neuroth, 1997; Ahrendt et al., 1998; Drost et al., 2003). The Ordovician sequence contains pyroclastic, effusive, and subvolcanic rocks (Štorch, 1998) that reach a thickness of up to 1000 m near the eruption centers (Fiala, 1971). Patocˇka et al. (1993) characterized the Lower and Upper Ordovician submarine effusives as alkaline basic to intermediate igneous rocks with within-plate signatures. In the Teplá-Barrandian unit the Late Ordovician glaciation in the Hirnantian is recorded by a glaciomarine diamictite at the base of the Kosov Formation and by a prominent glacio-eustatic regression resulting in storm-influenced sediments (Štorch, 1986, 1990; Brenchley and Štorch, 1989). The Ordovician-Silurian boundary is petrographically and paleontologically well documented. Upper Ordovician sediments of the Kosov Formation were replaced by Lower Silurian black graptolite shales of the Želkovice Formation (Havlícˇek and Vaneˇ k, 1966; Krˇ iž, 1998, and references therein). These organic-rich Lowermost Silurian “hot” shales form a prominent lithostratigraphic marker that can be traced across north Africa and the Arabian Peninsula (Lüning et al., 2000). A gap in sedimentation between uppermost Ordovician and upper Llandovery occurs only locally in the Teplá-Barrandian unit. SAMPLES AND METHODS Samples A total of twenty-two samples were taken to analyze the Nd-Sr-Pb isotopic composition of the Neoproterozoic basement and its Early Paleozoic overstep sequence. Rock types,
196
Drost et al.
lithostratigraphic units, and localities are listed in Table 1. The Neoproterozoic basement is represented by five graywacke samples and two shales. The Early Paleozoic succession is geochemically characterized on the basis of two Lower Cambrian fine-grained sandstones, one Lower Cambrian shale, five Middle Cambrian shales, one Lower to Middle Ordovician shale, five Upper Ordovician shales, and one Upper Ordovician siltstone. Major, Trace, and Rare Earth Element Data Whole-rock major, trace, and rare earth element (REE) contents were determined by ACTLABS (Activation Laboratories Ltd., Ancaster, Ontario, Canada) using lithium metaborate/tetraborate fusion and inductively coupled plasma mass spectrometry (ICP-MS), respectively. Geochemical data and detection limits are shown in Table 2. High quality of the data is ensured by regular analyses of certified reference materials and accredited through the International Organization for Standardization/International Electrotechnical Commission. Isotope Geochemistry Pb, Sr, and Nd were separated using standard ion exchange techniques. Procedures used at GeoForschungsZentrum Potsdam are described in Romer et al. (2001, 2005). Total proce-
dural blanks for whole-rock samples are 15–30 pg Pb, <100 pg Sr, and <50 pg Nd. Pb was loaded together with H3PO4 and silica-gel on single Re filaments (Gerstenberger and Haase, 1997). The isotopic composition of Pb was determined at 1200– 1250 °C on a Finnigan MAT262 multicollector mass spectrometer using dynamic multicollection. Instrumental fractionation was corrected with 0.1%/a.m.u., as determined from repeated measurement of lead reference material NBS 981. Accuracy and precision of the reported Pb ratios are better than 0.1% at the 2σ level. Sr was loaded on single Ta filaments, and its isotopic composition was determined on a VG 54–30 Sector multicollector mass spectrometer using a triple-jump dynamicmulticollection experiment. 87Sr/86Sr data are normalized with 86 Sr/88Sr = 0.1194. Repeated measurement of Sr standard NBS 987 during the measurement period gave 0.710249 ± 0.000004 (2σ, n = mean of twelve measurements). Nd was loaded on double Re filaments, and its isotopic composition was measured on a Finnigan MAT262 multicollector mass spectrometer using a double-jump dynamic-multicollection experiment. 143Nd/144Nd data are normalized with 146Nd/144Nd = 0.7219. Repeated measurement of La Jolla Nd standard during the measurement period gave 143Nd/144Nd = 0.511850 ± 0.000004 (2σ, n = mean of fourteen measurements). Analytical uncertainties of 87Sr/86Sr and 143Nd/144Nd are reported as 2σm. Isotopic compositions of the analyzed samples are given in Table 3.
TABLE 1. LIST OF ANALYZED SAMPLES Age
Lithostratigraphic unit
Locality
Coordinates* Y(west) X(south)
Sample
Rock type
Neoproterozoic 1 BL1 2 DB1/1 3 DB1/2 4 NPZb4 5 NPZb5 6 Dobris 7 Dobris 02B
Deformed graywacke Deformed graywacke Deformed graywacke Silty shale Shale with silty layers Fine-grained graywacke Fine-grained graywacke
Neoproterozoic Neoproterozoic Neoproterozoic Neoproterozoic Neoproterozoic Neoproterozoic Neoproterozoic
Blovice Formation Blovice Formation Blovice Formation Šteˇchovice Group Šteˇchovice Group Šteˇchovice Group Šteˇchovice Group
Vlcíce, Úslava River Týrovice Týrovice Prague-Zbraslav Prague-Zbraslav Dobríš Dobríš
813046 789327 789312 746698 746698 765875 765875
1091003 1047320 1047354 1058807 1058807 1075565 1075565
Cambrian 8 DB S1 9 DB D2 10 MM2 11 DB1/5 12 DB6/1 13 DB6/3 14 J4 15 J8
Fine-grained sandstone Fine-grained sandstone Shale Shale Shale Shale with silty layers Shale Silty shale
Lower Cambrian Lower Cambrian Lower Cambrian Middle Cambrian Middle Cambrian Middle Cambrian Middle Cambrian Middle Cambrian
Sádek Formation Sádek Formation Holšiny-Horice Formation Jince Formation Jince Formation Jince Formation Jince Formation Jince Formation
East of Bratkovice Northwest of Hluboš North of Hluboš North of Týrovice NNE of Skryje NNE of Skryje Jince Vinice Jince Vinice
777519 777788 777054 789291 790962 790964 777889 778512
1077124 1076180 1074725 1047482 1049548 1049422 1072233 1071645
Ordovician 16 KIRS 01 Shale Lower/Middle Ordovician Klabava Formation Rokycany strán 807683 1071834 17 LiBS0 Shale Upper Ordovician Liben Formation Prague, Bíla skála 740138 1040312 18 LeZb4 Silty shale Upper Ordovician Letná Formation Prague-Zbraslav 745989 1055390 19 ViBe 1A Shale Upper Ordovician Vinice Formation Beroun-North 769843 1052242 20 KDLe S2 Siltstone Upper Ordovician Králuv Dvur Formation Levín 774225 1056488 21 KoLe S1 Shale Upper Ordovician Kosov Formation Levín 774337 1056544 22 KoLe S2 Shale Upper Ordovician Kosov Formation Levín 774337 1056544 *Krovak coordinates, geodedic datum reference: S-JTSK (Czech Republic), fundamental station: Hermannskogel, ellipsoid: Bessel. Coordinates were read from 1:10,000 and 1:50,000 topographic maps; they do not have the accuracy of GPS (global positioning system) data.
Nd-Sr-Pb isotopic signatures of Neoproterozoic–Early Paleozoic siliciclastic rocks
TABLE 2. MAJOR, TRACE, AND RARE EARTH ELEMENT DATA DB1/1 DB1/2 NPZb 4 NPZb 5 Dobris Dobris 02B
197
Sample Detection (wt%) limit 0.01 SiO2 Al2O3 0.01 Fe2O3 0.01 MnO 0.001 MgO 0.01 CaO 0.01 0.01 Na2O K2O 0.01 0.001 TiO2 0.01 P2O5 LOI
71.1 13.2 3.94 0.051 1.55 0.50 4.02 2.02 0.513 0.15 1.81
65.9 15.7 5.25 0.097 2.17 0.40 3.95 2.43 0.618 0.17 3.14
Total
98.9
99.8
100.4
100.1
99.8
100.5
100.0
100.2
100.1
9 65 76 8 42 29 65 16 66 118 17 181 9 2.2 789 26.8 54.2 5.89 22.2 4.3 1.13 3.4 0.7 3.3 0.7 2.0 0.34 2.0 0.32 4.9 0.7 0.4 10 7.0 3.6
18 130 197 17 59 36 145 20 104 115 25 206 10 3.3 881 32.5 63.9 6.96 26.6 5.0 1.22 4.4 0.8 4.1 0.9 2.3 0.42 2.4 0.38 5.3 0.9 0.8 14 9.2 2.8
19 147 213 20 69 42 111 18 89 93 25 162 9 3.5 879 20.7 42.6 5.90 19.8 4.1 1.07 3.9 0.8 4.1 0.9 2.4 0.41 2.3 0.37 4.1 0.8 0.7 13 5.5 2.6
21 140 242 21 72 35 115 20 76 161 16 108 8 4.3 973 22.0 43.5 5.02 19.3 3.8 0.88 3.5 0.7 3.1 0.7 1.8 0.32 1.7 0.30 3.5 0.6 0.8 16 6.2 2.3
17 124 111 15 46 36 94 18 74 199 30 130 10 3.5 583 25.2 48.7 5.93 26.2 6.2 1.25 5.5 1.0 5.3 1.0 2.7 0.48 2.8 0.43 4.2 0.9 0.4 13 7.9 2.7
16 99 173 14 52 43 99 17 71 156 20 198 8 2.9 703 19.5 38.2 4.68 18.4 3.9 0.96 3.2 0.7 3.4 0.7 2.0 0.34 2.0 0.32 5.1 0.7 0.6 16 7.1 2.1
17 121 106 25 57 72 137 22 111 149 22 173 12 6.6 633 30.7 59.3 6.49 25.0 4.8 1.20 4.1 0.8 4.0 0.9 2.5 0.42 2.4 0.40 5.0 0.9 0.8 25 9.3 2.8
11 73 140 7 57 21 81 12 49 99 24 275 7 6.6 329 22.5 46.3 5.84 23.4 4.6 1.23 4.3 0.8 4.1 0.9 2.6 0.40 2.4 0.37 7.0 0.6 0.4 19 6.2 1.8
25 122 108 16 94 35 124 27 96 198 39 314 13 4.8 943 30.4 63.5 8.23 34.5 7.2 1.94 6.7 1.2 6.6 1.4 4.1 0.64 4.0 0.60 8.5 1.1 0.6 20 10.8 2.9
(ppm) Sc V Cr Co Ni Cu Zn Ga Rb Sr Y Zr Nb Cs Ba La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Hf Ta Tl Pb Th U
1 5 20 1 20 10 30 1 2 2 1 5 1 0.5 3 0.1 0.1 0.05 0.1 0.1 0.05 0.1 0.1 0.1 0.1 0.1 0.05 0.1 0.04 0.2 0.1 0.1 5 0.1 0.1
BL1
66.7 14.3 6.55 0.097 2.13 0.27 3.41 2.06 0.684 0.16 4.00
63.9 15.9 6.09 0.100 3.22 0.52 4.15 2.65 0.571 0.09 2.91
60.2 13.7 6.24 0.125 2.65 4.76 3.36 2.17 0.648 0.28 5.70
68.7 14.3 4.93 0.066 2.12 1.09 3.89 2.20 0.544 0.11 2.52
62.3 15.5 7.01 0.107 2.80 1.48 2.94 3.00 0.670 0.15 4.07
DB S1
DB D2
MM2
71.5 11.4 3.88 0.087 0.79 2.82 2.86 1.49 0.835 0.13 4.36
57.7 18.3 7.71 0.116 2.92 1.78 3.05 3.22 1.257 0.23 3.77
53.6 20.7 8.97 0.061 2.74 0.37 0.34 4.30 0.851 0.18 6.89 99.0 21 161 142 23 78 64 168 30 151 47 37 150 12 16.7 578 38.4 73.2 8.45 33.4 6.7 1.77 7.5 1.2 6.3 1.4 3.8 0.61 3.6 0.54 4.4 1.0 2.4 21 10.5 3.4 Continued
198
Sample (wt%) SiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O TiO2 P2O5 LOI Total
Drost et al.
TABLE 2. MAJOR, TRACE, AND RARE EARTH ELEMENT DATA (continued) DB6/3 J4 J8 KIRS 01 LiBS0 LeZb4 ViBe 1A KDLe S2
DB1/5
DB6/1
62.5 17.2 7.25 0.068 2.67 0.30 2.14 3.13 0.773 0.15 3.87
61.1 17.1 7.32 0.060 2.21 0.34 2.24 2.93 0.780 0.14 5.04
60.5 17.4 7.44 0.085 2.57 0.31 2.29 2.93 0.838 0.18 4.69
59.4 17.7 7.72 0.070 2.44 0.29 1.81 3.45 0.858 0.10 5.96
99.3
99.2
99.8
100.1
100.3
100.1
99.3
100.1
100.2
100.0
99.9
20 157 170 25 75 54 143 26 119 96 31 188 13 8.4 734 35.8 68.3 7.92 30.9 6.1 1.47 5.1 1.0 5.3 1.1 3.2 0.51 3.1 0.48 5.0 1.0 0.8 19 8.6 2.9
21 147 161 26 79 57 158 27 140 57 40 173 13 8.4 560 35.6 76.2 8.06 32.5 6.9 1.67 7.2 1.3 6.9 1.6 4.4 0.73 4.2 0.63 5.0 1.0 1.5 20 8.4 3.2
20 154 173 26 77 53 142 25 123 101 26 171 13 6.3 547 34.6 70.5 7.37 28.1 5.6 1.4 5.1 0.9 4.7 1.1 2.8 0.47 2.8 0.43 4.7 1.0 0.8 18 9 2.7
18 110 116 14 66 39 119 30 168 100 34 142 20 8.3 595 67.7 117 13.0 51.7 9.1 2.00 7.5 1.3 6.3 1.2 3.2 0.51 3.2 0.47 4.4 1.7 1.0 23 15.8 3.1
18 138 130 12 67 40 <30 33 175 118 23 184 20 8.1 667 58.0 111 12.6 46.0 6.7 1.36 5.2 0.8 3.9 0.9 2.8 0.47 2.6 0.43 5.5 1.6 0.5 10 16.9 3.2
26 160 164 8 77 26 <30 35 216 134 45 238 23 9.7 988 64.8 137 17.5 63.8 10.6 2.64 11.2 1.6 7.7 1.8 5.0 0.72 4.3 0.74 7.0 1.8 0.9 (10) 19.3 3.3
19 134 127 13 56 40 111 31 199 217 32 198 23 9.8 837 55.6 104 12.8 50.0 8.2 1.82 7.8 1.3 6.0 1.2 3.6 0.59 3.4 0.52 6.4 1.7 1.4 22 15.6 2.9
21 179 145 52 88 44 94 28 154 134 24 235 22 8.9 412 43.0 93.6 10.4 37.5 5.6 1.24 4.8 0.8 4.0 1.0 3.0 0.49 2.9 0.51 7.1 1.6 0.7 18 13.1 2.6
21 155 142 29 63 43 121 30 146 127 27 190 19 7.8 385 45.2 97.4 10.9 38.8 6.0 1.42 5.3 0.9 4.3 1.1 3.1 0.48 2.9 0.51 5.7 1.4 0.9 20 13.1 2.7
17 153 143 27 70 37 137 24 135 108 24 247 20 6.5 367 42.6 78.7 8.48 34.0 5.1 1.04 4.1 0.8 4.1 0.9 2.8 0.50 2.9 0.45 7.5 1.5 0.9 26 11.2 2.5
100.1
(ppm) Sc 18 19 V 152 146 Cr 124 141 Co 22 20 Ni 58 77 Cu 44 48 Zn 130 149 Ga 22 24 Rb 136 108 Sr 78 96 Y 25 28 Zr 188 167 Nb 11 11 Cs 5.9 5.7 Ba 667 695 La 28.0 30.4 Ce 57.6 58.1 Pr 6.42 7.15 Nd 25.0 28.5 Sm 5.1 5.8 Eu 1.17 1.44 Gd 4.3 5.1 Tb 0.8 0.9 Dy 4.3 5.2 Ho 0.9 1.1 Er 2.6 2.9 Tm 0.46 0.49 Yb 2.7 3.0 Lu 0.44 0.45 Hf 5.0 4.5 Ta 0.9 1.0 Tl 0.7 0.7 Pb 17 14 Th 7.6 7.5 U 2.9 2.5 Note: LOI—loss on ignition.
61.4 16.8 7.27 0.056 2.88 0.34 2.79 2.95 0.937 0.19 4.53
58.1 22.2 6.41 0.012 1.58 0.27 0.53 3.43 0.992 0.16 6.61
56.0 23.8 5.39 0.018 1.32 0.15 0.62 3.61 1.03 0.09 8.07
54.3 24.4 3.90 0.015 1.58 0.27 0.52 6.11 1.38 0.19 6.63
59.4 21.2 4.68 0.086 0.88 0.32 0.68 4.67 1.17 0.25 6.79
61.9 18.5 5.72 0.104 1.92 0.74 0.88 3.42 1.26 0.11 5.6
KoLe S1
KoLe S2
58.6 20.5 6.92 0.060 1.94 0.41 0.65 3.47 1.11 0.14 6.2
65.0 16.2 6.41 0.094 1.64 0.32 0.87 3.16 1.06 0.12 4.98
†
87
86
§
87
86
#
TABLE 3. WHOLE-ROCK Sr, Nd, AND Pb ISOTOPE DATA 143
144
§
206
204
††
207
204
††
208
204
††
206
204
§§
207
204
§§
208
204
§§
# TDM** Sr/ Sr SrT/ Sr Nd/ Nd Pb/ Pb Pb/ Pb Pb/ Pb PbT/ Pb PbT/ Pb PbT/ Pb T εNdT (Ma) (Ga) 1 BL1 570 0.721829 ± 7 0.7087 0.511733 ± 5 –11.9 2.17 18.152 15.519 38.483 16.05 15.40 37.18 2 DB1/1 570 0.723299 ± 7 0.7020 0.511978 ± 4 –6.8 1.80 19.475 15.628 39.210 18.19 15.55 37.97 3 DB1/2 570 0.723890 ± 7 0.7014 0.512182 ± 5 –3.8 1.49 20.094 15.666 38.721 18.70 15.58 37.83 4 NPZb4 570 0.717642 ± 10 0.7065 0.512069 ± 6 –5.5 1.66 18.597 15.599 38.541 17.75 15.55 37.81 5 NPZb5 570 0.718628 ± 7 0.7098 0.512105 ± 4 –6.5 1.61 18.895 15.616 38.911 17.66 15.54 37.76 6 Dobris 570 0.718490 ± 7 0.7078 0.512015 ± 5 –7.2 1.74 18.516 15.603 38.662 17.78 15.56 37.84 7 Dobris 02B 570 0.723902 ± 10 0.7064 0.512041 ± 12 –5.7 1.70 18.332 15.587 38.497 17.68 15.55 37.80 8 DB S1 520 0.714348 ± 8 0.7037 0.512425 ± 6 1.0 1.12 18.933 15.672 38.800 18.36 15.64 38.25 9 DB D2 520 0.715030 ± 7 0.7046 0.512414 ± 4 0.3 1.13 19.003 15.678 38.863 18.19 15.63 37.91 10 MM2 519 0.770952 ± 7 0.7018 0.512193 ± 5 –3.7 1.47 19.352 15.708 39.225 18.57 15.66 38.39 11 DB1/5 505 0.736003 ± 18 0.6996 0.512140 ± 5 –5.0 1.55 18.806 15.618 38.679 17.88 15.57 37.88 12 DB6/1 505 0.732287 ± 9 0.7088 0.512162 ± 5 –4.5 1.52 18.857 15.613 38.739 18.10 15.57 37.89 13 DB6/3 505 0.734657 ± 13 0.7088 0.512076 ± 4 –6.0 1.65 18.567 15.595 38.510 17.74 15.55 37.71 14 J4 505 0.753715 ± 10 0.7023 0.512323 ± 5 –1.7 1.27 19.150 15.626 38.728 18.36 15.58 38.05 15 J8 505 0.731392 ± 10 0.7060 0.512079 ± 6 –6.0 1.65 18.811 15.610 38.862 17.93 15.56 38.02 16 KlRS 01 470 0.744340 ± 12 0.7117 0.511876 ± 5 –9.6 1.96 18.561 15.695 39.195 17.91 15.66 38.12 17 LiBS0 455 0.741838 ± 14 0.7139 0.511834 ± 4 –9.4 2.02 18.580 15.693 39.198 17.07 15.61 36.64 18 LeZb4 455 0.738781 ± 18 0.7083 0.511928 ± 4 –8.3 1.88 18.842 15.708 39.654 17.27 15.62 36.70 19 ViBe 1A 455 0.727928 ± 13 0.7107 0.511902 ± 5 –8.7 1.92 18.776 15.711 39.282 18.15 15.68 38.20 20 KDLe S2 445 0.730318 ± 14 0.7091 0.511869 ± 6 –9.0 1.97 18.723 15.702 39.094 18.06 15.66 38.02 21 KoLe S1 445 0.730996 ± 10 0.7099 0.511883 ± 6 –8.9 1.95 18.706 15.688 39.083 18.08 15.65 38.11 22 KoLe S2 445 0.734525 ± 7 0.7114 0.511891 ± 5 –8.6 1.93 18.618 15.681 38.915 18.18 15.66 38.28 *Samples were dissolved with 52% HF for four days at 160 °C on the hot plate. Digested samples were dried and taken up in 6N HCl. Sr and Nd were separated and purified using ion-exchange chromatography as described in Romer et al. (2001). Pb was separated using the HBr-HCl ion-exchange procedure of Tilton (1973) and Manhès et al. (1978). † Stratigraphic ages were estimated using available biostratigraphic and geochronological data and the time scale of Gradstein et al. (2005). § 87 86 143 144 86 88 146 144 Sr/ Sr and Nd/ Nd are normalized to Sr/ Sr = 0.1194 and Nd/ Nd = 0.7219, respectively. Sr data were obtained on VG 54-30 multicollector mass spectrometer using a triple-jump dynamic-multicollection experiment. Nd data were obtained on a Finnigan MAT262 multi-collector mass spectrometer using a double-jump dynamic multicollection experiment. Analytical uncertainties are given at 2σm level. 86 87 –11 0 # 87 –1 147 –12 –1 147 144 143 144 0 Sr/ SrT and εNdT were calculated for the stratigraphic age using λ Rb = 1.42 × 10 y and λ Sm = 6.54 × 10 y , ( Sm/ Nd) CHUR = 0.1967, and ( Nd/ Nd) CHUR = 0.512638, respectively, and the concentration data given in Table 2. **Crustal residence ages were calculated with the parameters given by Liew and Hofmann (1988). †† Lead isotope data corrected for mass discrimination with 0.1%/a.m.u. Reproducibility at 2σ level is better than 0.1%. §§ Lead isotope data recalculated to the biostratigraphic age using the contents of Pb, Th, and U (Table 2) and the constants of Jaffey et al. (1971) recommended by the International Union of Geological Sciences (Steiger and Jäger, 1977).
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RESULTS Elemental Geochemistry Analyzed samples from Neoproterozoic to Ordovician siliciclastic sequences of the Teplá-Barrandian unit (Table 2) vary in SiO2 content from 53.6 to 71.6 wt%. Al2O3/SiO2, K2O/Na2O, and Al2O3/(CaO + Na2O) ratios increase from Neoproterozoic to Ordovician samples, reflecting increasing weathering and decomposition of detrital minerals, particularly plagioclase, as well as increasing clay mineral content (Fig. 3A). Lower Cambrian samples differ from the general trend, whereas high values of shale sample MM2 are in the same range as Ordovician rocks displaying intense weathering after the Cadomian orogeny. Low values of finegrained sandstones overlap with those of the Neoproterozoic and/ or Middle Cambrian sediments, respectively. Similarly the chemical index of alteration (CIA; Nesbitt and Young, 1982) increases from Neoproterozoic (46–64) to Ordovician (77–96) sedimentary rocks. The CIA of the Lower Cambrian siliciclastics varies from 51 to 79, whereas the shale sample MM2 displays the highest value, and fine-grained sandstones (51, 62) fall in the same range as Neoproterozoic rocks. Middle Cambrian siliciclastics have CIA values of 71–78 that are markedly higher than those of Neoproterozoic samples, but lower than those of Ordovician rocks (77–96). Post-Archean average Australian shale (PAAS)-normalized multi-element diagrams (Fig. 4A–C) reveal that Cs, Rb, Th, K, and light rare earth element (LREE) contents increase from Neoproterozoic to Ordovician sediments. As these elements are enriched in the earth’s crust when compared to the mantle, this change suggests an increasing contribution of crustal sources toward the Ordovician. The progressive enrichment of these incompatible elements is controlled by reworking of a weathering crust during the Lower Cambrian and most of the Ordovician. Ca (not shown) and Sr, which are mobile during weathering processes (Nesbitt et al., 1980), are depleted in the Cambrian and Ordovician samples relative to PAAS. Lower Cambrian shale sample MM2 differs from the general trend, having Cs and Rb concentrations in the same range as PAAS. Although the alkali and alkaline earth elements may be subject to redistributions during diagenesis and low-grade metamorphism (Taylor and McLennan, 1985; Wronkiewicz and Condie, 1987), the change of contents and ratios in the analyzed samples with time suggests different intensities of source-rock weathering and changing provenances (cf. Nesbitt et al., 1980) rather than synand postdepositional alterations, which is supported by generally parallel Cs-Rb patterns in the spider diagrams (Fig. 4A–C). PAAS-normalized REE data of Neoproterozoic (Fig. 4D) and Middle Cambrian (Fig. 4E) siliciclastics show a depletion of LREE relative to the heavy REE (HREE) and a positive Eu anomaly. Lower Cambrian sandstone samples are somewhat more strongly depleted in LREE. REE patterns of Ordovician pelites largely resemble PAAS (Fig. 4F), as is typical for sediments that were mixed efficiently during derivation from a large area of stable continental crust.
There is no significant correlation of any geochemical feature with the grain size of the samples. Instead, geochemical features strongly correlate with the depositional age of the sedimentary rocks. Therefore the variation in geochemical compositions is attributed to provenance and weathering conditions rather than a result of sedimentary sorting, transport, or dilution by quartz. Isotope Geochemistry Six of seven Neoproterozoic samples yield εNdT values (T = age of deposition) of −7.2 to −3.8 and TDM from 1.49 to 1.80 Ga (Table 3; Figs. 5 and 6). Sample BL1 exhibits the strongest negative εNd570 value of −11.9 and the oldest TDM (2.17 Ga). All TDM values were calculated using the two-step model of Liew and Hofmann (1988). 87Sr/86Sr570 ratios of five samples range between 0.7064 and 0.7098; the remaining two samples (DB1/1, DB1/2) reveal anomalously low 87Sr/86Sr570 ratios resulting from Rb gain or Sr loss. Rb/Sr ratios of these samples (0.90, 0.96) are slightly higher than those of the other Precambrian samples (0.37–0.74) and caused an overcorrection of the in situ formed 87Sr. The Sr evolution lines of the apparently undisturbed Neoproterozoic sediments define a narrow field (Fig. 5D). Intermediate LaN/YbN ratios (6.0–9.3; norm: chondrite; Boynton, 1984) of the Neoproterozoic siliciclastic sediments are in agreement with the Nd and Sr isotopic data, demonstrating mixing of a juvenile component derived from the Cadomian magmatic rocks and older continental crust (Fig. 5A). Cambrian samples have higher εNdT values (−6.0 to +1.0) than do the Neoproterozoic ones, even attaining positive values in two Lower Cambrian samples (DB S1, DB D2), and have younger TDM (1.65–1.12 Ga). 87Sr/86SrT ratios of five samples vary from 0.7037 to 0.7088; three analyzed shales (MM2, DB1/5, J4), however, have obviously disturbed Rb/Sr systematics, as the anomalously low 87Sr/86SrT ratios of 0.6996–0.7023 indicate overcorrection of in situ growth of 87Sr. REE characteristics, particularly of the Middle Cambrian sedimentary rocks, are similar to those of the analyzed Neoproterozoic rocks (Figs. 4 and 5A). Although major and trace element geochemical data suggest a derivation of the Cambrian siliciclastics predominantly from the Cadomian basement, the Nd and Sr isotopic signatures demonstrate a contribution of young mantle-derived material to these rocks (Table 3; Figs. 5 and 6). Two Lower Cambrian fine-grained sandstones have the youngest TDM (1.13–1.12 Ga; Fig. 6). Ordovician pelites exhibit very homogenous εNdT values of −9.6 to −8.3 and TDM of 2.02–1.88 Ga (Figs. 5 and 6). The latter overlap with the old crustal residence ages of Neoproterozoic sediments. The high TDM and the high 87Sr/86SrT ratios, varying from 0.7083 to 0.7139, indicate a significant proportion of old continental crust in the source area. REE patterns resemble those of PAAS, demonstrating a broad input of differentiated continental crust as well. Initial uranogenic lead isotopic compositions of Neoproterozoic and Middle Cambrian sediments fall in the same field (i.e., reflect sediment derivation from reservoirs of similar evolution;
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Figure 4. PAAS-normalized multi-element diagrams and REE patterns of Neoproterozoic to Ordovician siliciclastic rocks of the Barrandian. Normalization values from Taylor and McLennan (1985). (B) and (E): White circles stand for Lower Cambrian samples, black circles for Middle Cambrian siliciclastic rocks.
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Figure 5. Nd and Sr isotopic characterization of Neoproterozoic and Early Paleozoic siliciclastic sediments of the Barrandian. (A) εNdT vs. LaN/YbN diagram (normalized against chondrite composition as given by Boynton, 1984) showing the variable portion of mantle-derived and crustal material in the sediments. Note the particularly high portion of mantle-derived material in the Lower Cambrian sediments (high εNdT and low LaN/YbN values) and the predominance of crustal material in the Ordovician sediments (low εNdT and high LaN/YbN values). (B) εNdT vs. 87Sr/86SrT diagram illustrating varying contributions of old crustal material to the Neoproterozoic, Cambrian, and Ordovician siliciclastics. Samples in shaded area have anomalously low 87Sr/86SrT ratios that reflect an open Rb/Sr system, which allowed either late Rb addition or late Sr loss. (C) εNdT vs. age diagram highlighting the heterogeneous Nd isotopic signature of Cambrian sedimentary rocks and the uniform, low εNdT values of the Ordovician pelites. (D) Sr evolution diagram. Stippled field comprises the Sr evolution curves of the Ordovician samples.
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Nd-Sr-Pb isotopic signatures of Neoproterozoic–Early Paleozoic siliciclastic rocks Fig. 7A). They have, however, markedly lower 207Pb/204Pb values than do the Ordovician samples. Because most 207Pb formed early in the Earth’s history, when most of its parent 235U was still present, high 207Pb/204Pb values for a given 206Pb/204Pb indicate an older age for the 238U/204Pb increase in the Pb source. Because the formation of continental crust is generally associated with an increase in 238U/204Pb, the lower 207Pb/204Pb values in the Neoproterozoic and Middle Cambrian sediments indicate that they were derived from a younger crustal source. Thus, 207Pb/204PbT ratios of Lower Cambrian and Ordovician samples point to higher proportions of old continental crust in the source area. The 206Pb/204PbT versus 208Pb/204PbT diagram (Fig. 7B) distinguishes different reservoirs with distinct initial Th/U ratios. Neoproterozoic and Middle Cambrian samples reveal similar time-integrated Th/U ratios of ~3.1, which is compatible with derivation from magmatic arc rocks. Lower Cambrian siliciclastics had somewhat higher Th contents (Th/U ~3.4) that can be explained by additional involvement of old sedimentary rocks, as suggested by the isotopic composition of the uranogenic lead. In contrast to Neoproterozoic and Cambrian samples, Ordovician pelites have distinctly higher initial Th/U ratios of ~5.2. DISCUSSION AND CONCLUSIONS Variations of Isotopic Signatures and Disturbance of Isotopic Systems The isotope data are in agreement with major and trace element compositions and clearly vary among samples of different stratigraphic ages. However, in spite of the well-defined groups corresponding to different provenances during individual stratigraphic periods, there is some variation within these groups that
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has to be attributed to a combination of initial isotopic heterogeneities and under- or overcorrection for in situ growth. Two Ordovician samples (LiBS0, LeZb4) form outliers in the 206Pb/204PbT versus 207Pb/204PbT and 206Pb/204PbT versus 208Pb/204PbT plots. Their anomalous positions in the diagrams reflect overcorrection of in situ Pb growth and thus indicate disturbance of the U-Th-Pb system by lead loss. In contrast, the isotopic diversity of the Neoproterozoic samples cannot solely be attributed to alteration. The range in Pb and Nd isotopic compositions within Neoproterozoic rocks is mainly due to the samples of the Blovice Formation (lower part of the Neoproterozoic succession; Table 1; Fig. 2), whereas samples from the Šteˇ chovice Group (upper part of the Neoproterozoic succession; Table 1; Fig. 2) cluster in a very close range. Because the Blovice Formation was affected by a late Neoproterozoic metamorphic event (Pašava and Amov, 1993), the mobility of U6+ during metamorphism would be a suitable reason for disturbance of the U-Th-Pb system in the analyzed deformed graywackes. Thus, the lead isotopic data of the samples from the Blovice Formation (BL1, DB1/2) do not necessarily reflect the provenance of these rocks but are probably the result of U redistribution during late Neoproterozoic metamorphism. However, the variations in the Nd isotope compositions of the Neoproterozoic samples indicate changing proportions of material derived from different sources. The larger the input from Neoproterozoic magmatic arc rocks and the lower the contribution from old crustal sources, the more radiogenic is the Nd isotope signature. Sample BL1 from the Blovice Formation differs from the younger Neoproterozoic rocks with respect to the U-Th-Pb and Sm-Nd systems. Although its 87Sr/86Sr570 ratio of 0.7087 is a reasonable value for detrital material from an old craton, it cannot be excluded that this sample also underwent Rb gain or Sr loss.
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Figure 7. (A) 206Pb/204Pb(T) vs. 207Pb/204Pb(T) diagram distinguishing between different reservoirs of uranogenic lead. (B) 206Pb/204Pb(T) vs. 208Pb/ 204 Pb(T) diagram distinguishing varying initial Th/U ratios related to different sources. Note that the anomalous unradiogenic position of samples BL1, LiBS0, and LeZb4 may reflect late Pb loss, which results in an overcorrection of in situ Pb growth.
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The combined Nd-Sr isotopic data of sample BL1 indicate major input from old high-grade metamorphic crust (unradiogenic εNdT at low 87Sr/86SrT ratios). The heavy mineral spectrum and the petrography of the sample, however, rather suggest granitoid detritus. If the 87Sr/86Sr570 ratio of 0.7087 reflects overcorrection of in situ Sr growth, the geochemical and isotopic data as well as the petrographic features may be explained by major input from Paleoproterozoic granitoid rocks, such as the Icartian basement in the Armorican Massif. The Rb-Sr system of Neoproterozoic samples DB1/1 and DB1/2, as well as of the Cambrian shales MM2, DB1/5, and J4, is disturbed. These samples yielded geologically unreasonably low 87Sr/86SrT ratios (<0.704) because of overcorrection of in situ 87Sr growth. For instance, late Rb addition or Sr loss because of the relatively high mobility of Rb and Sr during fluidrock interaction at low to medium temperatures would increase Rb/Sr ratios and eventually result in anomalously low apparent 87 Sr/86SrT ratios. Such a fractionation of Rb and Sr is reflected in the Sr evolution trends, which are distinctly steeper than in apparently undisturbed samples (Fig. 5D). Crustal Residence Ages Single-stage TDM ages (e.g., DePaolo, 1981) of siliciclastic sedimentary rocks reflect the average crustal residence time
of all contributing sources (i.e., the mixing of material that was derived from the depleted mantle at different times). Two-stage TDM ages, as presented by Liew and Hofmann (1988), assume that old crust and juvenile rocks were sediment sources and that the 147Sm/144Nd value of the cratonic source is not identical with the measured one, which represents a weighted mixture of both sources. The average crustal residence time of the crustal source is estimated by using a 147Sm/144Nd ratio of 0.12, which is a representative value for average cratonic crust. All our new TDM ages are two-stage model ages. Teplá-Barrandian Neoproterozoic to Ordovician siliciclastic rocks gave TDM values of 2.17–1.12 Ga, whereas seventeen of the twenty-two samples yield crustal residence ages >1.5 Ga. Figure 8 shows the Nd isotopic evolution of Neoproterozoic–Early Paleozoic (this study) and Devonian (Strnad and Mihaljevicˇ, 2005) siliciclastic sedimentary rocks of the Teplá-Barrandian unit in comparison to potential source rocks. Besides Paleoproterozoic basement rocks known to be present in the northern Armorican Massif (Auvray et al., 1980; Samson and D’Lemos, 1998; Inglis et al., 2004) and the Moldanubian domain in the southern Bohemian Massif (Wendt et al., 1993; Friedl et al., 2004), the Cadomian magmatic arc supplied detrital material to the Neoproterozoic siliciclastic rocks of the Teplá-Barrandian unit (Fig. 8A). The Nd isotopic signature of sample BL1 is similar to that of the Icartian basement of the Armorican Massif. The other six Neoproterozoic samples contain more prominent proportions
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Age (Ga) Figure 8. εNd evolution diagrams comparing the analyzed samples and potential source rocks. (A) Squares represent εNd570 values for Neoproterozoic siliciclastic sedimentary rocks of the Teplá-Barrandian unit. (B) Large symbols represent time-integrated εNd values of Cambrian to Devonian Teplá-Barrandian sedimentary rocks (triangles = εNd values of Devonian graywackes from Strnad and Mihaljevicˇ, 2005). Depleted mantle curve from Liew and Hofmann (1988). CHUR—chondritic uniform reservoir; TBU—Teplá-Barrandian unit. Data sources: *1—Pin and Waldhausrová (this volume); *2—Vokurka and Frýda (1997); *3—Pin et al. (this volume); *4—Liew and Hofmann (1988); *5—Wendt et al. (1993); *6—Samson and D’Lemos (1998).
Nd-Sr-Pb isotopic signatures of Neoproterozoic–Early Paleozoic siliciclastic rocks of juvenile magmatic arc–derived material. Nevertheless, their TDM values (≥1.5 Ga), which are distinctly older than the deposition age, emphasize the involvement of old crustal sources. The Cambrian samples have younger TDM (1.65–1.12 Ga) than do the Neoproterozoic sediments, which is compatible with major input from synsedimentary mantle-derived volcanism related to incipient rifting (Fig. 8B). Although Ordovician to Devonian strata of the Teplá-Barrandian unit also contain volcanic rocks with highly radiogenic Nd isotopic signatures (Vokurka and Frýda, 1997; Pin et al., this volume; Fig. 8B), the sedimentary record of this period seems to be less influenced by the juvenile source. Ordovician detrital sediments have TDM values of 2.02– 1.88 Ga, which are distinctly older than those of the Cambrian sedimentary rocks (Fig. 8B) and reflect major contributions from old crustal sources. Devonian graywackes studied by Strnad and Mihaljevicˇ (2005) have TDM values of 1.75–1.58 Ga (recalculated to the two-stage model of Liew and Hofmann, 1988) that overlap those of the Neoproterozoic graywackes. Provenance and Geotectonic Setting Both the element geochemical and isotope data of the analyzed rocks clearly form individual groups of samples. These groups match the stratigraphic ages of the analyzed siliciclastic rocks and reflect changing sources and weathering conditions over time, which in turn indicate shifts in the geotectonic setting. Neoproterozoic siliciclastic rocks were deposited in an active geotectonic setting (Fig. 9). Tectonic activity ensured fast supply of fresh material to the basin, which is reflected in the major and trace element compositions of the Neoproterozoic sedimentary rocks. Although detritus derived from the Cadomian magmatic arc dominates the chemical and petrographic compositions of the Late Neoproterozoic sediments (cf. Jakeš et al., 1979; Lang, 2000), Nd and Sr isotopic signatures reveal mixing of juvenile material with detritus from an old crustal source (Figs. 5 and 8A). Particularly, the fact that the TDM values (2.17–1.49 Ga) are significantly older than the Late Neoproterozoic stratigraphic age points to involvement of differing proportions of old continental crust. This observation is in agreement with findings of Proterozoic to Archean detrital and inherited zircon in graywackes and granitoid pebbles from conglomerates of the Neoproterozoic succession (Dörr et al., 2002; Sláma et al., 2003; Drost et al., 2004). Detrital zircon from the Šteˇ chovice Group indicates Neoproterozoic sedimentation until at least ca. 570 Ma (Dörr et al., 2002; Drost et al., 2004). Intrusions emplaced within the Cadomian basement of the Teplá-Barrandian unit during transtensional movements at 524 Ma and later postdate the deposition of Cadomian volcanosedimentary successions and indicate that the change of the geotectonic regime from subduction to transtension/rifting was initiated before 524 Ma (Zulauf et al., 1997b; Dörr et al., 1998, 2002). Lower Cambrian, mainly continental sediments have more radiogenic εNdT values (−3.7 to +1.0) and younger TDM (1.47– 1.12 Ga) than do the Neoproterozoic rocks. Initial 87Sr/86Sr ratios
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(~0.704), too, mirror significant input of detritus from young mantle-derived magmatic rocks. Juvenile material was provided by synsedimentary volcanism, causing the relatively unradiogenic initial Sr isotopic compositions of the Lower Cambrian sediments. In contrast, the uranogenic lead shows that old continental crust contributed significantly to these sedimentary rocks, that is, the Lower Cambrian physiography permitted access of craton-derived detritus to the Prˇ íbram-Jince basin. The contrasting indication of predominant sediment source obtained from Sr and Pb isotopes is not in conflict. Instead, it reflects the different contents of Sr and Pb in mantle-derived volcanic rocks and continental crust, respectively. The mantle-derived rocks have higher Sr contents and relatively low Pb contents, whereas the continental crust has lower Sr and relatively high Pb contents. In sediments containing contributions from both reservoirs, the Sr isotopic composition is dominated by the mantle-derived rocks, whereas the Pb isotopic signature is controlled by the crustal source. Deposition in a transtensional setting is in agreement with the geochemical data (Fig. 9) and furthermore supported by structural evidence from Cambrian granitoids in the western part of the Teplá-Barrandian unit that were emplaced synkinematically with respect to transtensive shearing (Zulauf et al., 1997b). Rift volcanism starting in Lower Cambrian times is also known from other crustal fragments with Cadomian basement and Early Paleozoic overstep sequences, such as the Torgau-Doberlug synform at the northern margin of the Bohemian Massif (Jonas et al., 2000) or the northeastern units of the Ossa-Morena zone in the Iberian Massif (Sánchez-García et al., 2003).
100 Ordovician Middle Cambrian Lower Cambrian
PM
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10
ACM 1
ARC
0.1 45
50
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65
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SiO2 (wt%) Figure 9. SiO2 vs. K2O/Na2O discrimination diagram (after Roser and Korsch, 1986) indicating an active geotectonic regime for the Neoproterozoic and Cambrian siliciclastics (except sample MM2) and a passive margin setting for the Ordovician pelites. ACM—active continental margin (including strike-slip margins); ARC—oceanic island arc; PM—passive continental margin.
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A major transgression led to the deposition of a Middle Cambrian marine sedimentary sequence with geochemical and isotopic characteristics largely corresponding to those of the Neoproterozoic sedimentary rocks. However, the markedly younger TDM values (1.65–1.27 Ga) reveal contributions from post-Neoproterozoic magmatic rocks as a new component. Indications for a provenance from the craton have not been detected in the Middle Cambrian siliciclastic rocks. Marine sedimentation was interrupted during the Upper Cambrian, when ~1500 m of subaerial volcanics (Waldhausrová, 1971) of lower crustal or mantle origin (Vidal et al., 1975) were deposited. During this time, advanced crustal thinning enabled asthenospheric updoming that was associated with uplift and volcanic activity. Subsequently, incipient thermal subsidence gave rise to a major transgression during the Tremadocian. The overlying Ordovician succession is completely marine. The sampled Ordovician pelites form a very homogeneous group with respect to their geochemical and isotopic compositions. Major and trace element data indicate reworking of an old weathered crust (cf. Patocˇka and Štorch, 2004). εNdT values of −9.6 to −8.3, as well as Sr and Pb isotope data of Ordovician pelites, reflect an increased input of material delivered from the Gondwana hinterland. TDM values range from 2.02 to 1.88 Ga. The uniform geochemical and isotopic signatures of the Ordovician samples point to a mature rift or shelf environment (Fig. 9). Sediments were derived from a large area of stable continental crust and were mixed efficiently to average out geochemical and isotopic heterogeneities of the source area. Givetian graywackes represent the youngest pre-Variscan rocks of the Teplá-Barrandian unit. Geochemical and Nd isotope data as well as detrital zircon ages of these rocks were published by Strnad and Mihaljevicˇ (2005). Geochemical and Nd isotope data (εNd380 values: −8.1 to −6.8) revealed significant contents of differentiated crustal material (cf. Fig. 8B). This result is substantiated by the age spectrum of detrital zircon containing numerous Paleoproterozoic and Archean grains. Furthermore the detrital zircon ages indicate a Gondwanan (north African) rather than Baltic provenance for these sediments (Strnad and Mihaljevicˇ, 2005). The geochemical and isotopic signatures of Neoproterozoic to Ordovician siliciclastic sedimentary rocks of the TepláBarrandian unit show an almost continuous input from an old crustal source that was diluted, particularly during the Cambrian, by Cadomian and post-Cadomian detritus derived from juvenile mantle-derived protoliths. In the Middle Cambrian crustal tilting starved the rift/wrench basin of sediments from the cratonic hinterland. The progressive enrichment of incompatible elements and the increase of source-rock weathering from the Neoproterozoic through the Cambrian to the Ordovician also support the continuous availability of material from the same, but increasingly weathered, hinterland. Therefore, after the separation of Avalonia, the Teplá-Barrandian unit should have represented—at least until the end of the Ordovician, but probably during the entire pre-Variscan Paleozoic—a part of the southern margin of the Rheic Ocean.
ACKNOWLEDGMENTS This study was supported by the Deutsche Forschungsgemeinschaft grant Li 521/14-1/2. A. Gerdes, C. Pin, and P. Budil are thanked for discussions. The manuscript has benefited from the constructive reviews of J. Hladil and M. Timmerman. This chapter is a contribution to the International Geological Correlation Program Project 497, “The Rheic Ocean: Its origin, evolution and correlatives.” REFERENCES CITED Ahrendt, H., Wemmer, K., and Neuroth, H., 1998, K-Ar-systematics on detrital white micas and fine mineral fractions from the Barrandian of the Prague syncline/Czech Republic: Acta Universitatis Carolinae—Geologica, v. 42, p. 204. Alvaro, J.J., Elicki, O., Geyer, G., Rushton, A., and Shergold, J., 2003, Palaeogeographical controls on the Cambrian trilobite immigration and evolutionary patterns reported in the western Gondwana margin: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 195, p. 5–35, doi: 10.1016/ S0031-0182(03)00300-6. Auvray, B., Charlot, R., and Vidal, P., 1980, Donnés nouvelles sur le Proterozoïque inférieur du domaine nord-Armoricain (France): âge et signification: Canadian Journal of Earth Sciences, v. 17, p. 532–538. Beetsma, J.J., 1995, The Late Proterozoic/Paleozoic and Hercynian crustal evolution of the Iberian Massif, N Portugal as traced by geochemistry and Sr-Nd-Pb isotope systematics of pre-Hercynian terrigenous sediments and Hercynian granitoids [Ph.D. thesis]: Amsterdam, Vrije Universiteit Amsterdam, 223 p. Boynton, W.V., 1984, Geochemistry of the rare earth elements: Meteorite studies, in Henderson, P., ed., Rare earth element geochemistry: Amsterdam, Elsevier, p. 63–114. Brenchley, P.J., and Štorch, P., 1989, Environmental changes in the Hirnantian (Upper Ordovician) of the Prague basin, Czechoslovakia: Geological Journal, v. 24, p. 165–181. Cháb, J., 1993, General problems of the TB (Teplá-Barrandian) Precambrian, Bohemian Massif, The Czech Republic: Veˇstník Cˇeského geologického ústavu, v. 68, p. 1–6. Cháb, J., Šrámek, J., Pokorný, L., Chlupácˇová, M., Manová, M., Vejnar, Z., Waldhausrová, J., and Žácˇek, V., 1997, C.4, The Teplá-Barrandian unit, in Vrána, S., and Šteˇdrá, V., eds., Geological model of western Bohemia related to the KTB borehole in Germany: Journal of Geological Sciences, v. 47, pp. 80–82. Chaloupský, J., Chlupácˇ, I., Mašek, J., Waldhausrová, J., and Cháb, J., 1995, VII.B.1, Stratigraphy, in Dallmeyer, R.D., Franke, W., and Weber, K., eds., Pre-Permian geology of central and eastern Europe: Berlin, Heidelberg, New York, Springer, p. 379–391. Chlupácˇ, I., 1993, Geology of the Barrandian—A field trip guide. SenckenbergBuch 69: Frankfurt am Main, Verlag W. Kramer, 163 p. Chlupácˇ, I., 1995, Lower Cambrian arthropods from the Paseky shale (Barrandian area, Czech Republic): Journal of the Czech Geological Society, v. 40, p. 9–36. Chlupácˇ, I., Havlícˇek, V., Krˇiž, J., Kukal, Z., and Štorch, P., 1998, Palaeozoic of the Barrandian (Cambrian to Devonian): Prague, Czech Geological Survey, 183 p. DePaolo, D.J., 1981, Neodymium isotopes in the Colorado front Range and crust-mantle evolution in the Proterozoic: Nature, v. 291, p. 193–196, doi: 10.1038/291193a0. Dörr, W., Fiala, J., Vejnar, Z., and Zulauf, G., 1998, U-Pb zircon ages and structural development of metagranitoids of the Teplá crystalline complex: Evidence for pervasive Cambrian plutonism within the Bohemian Massif (Czech Republic): Geologische Rundschau, v. 87, p. 135–149, doi: 10.1007/s005310050195. Dörr, W., Zulauf, G., Fiala, J., Franke, W., and Vejnar, Z., 2002, Neoproterozoic to Early Cambrian history of an active plate margin in the Teplá-Barrandian unit—A correlation of U-Pb isotopic-dilution-TIMS ages (Bohemia, Czech Republic): Tectonophysics, v. 352, p. 65–85, doi: 10.1016/S00401951(02)00189-0. Drost, K., Linnemann, U., Wemmer, K., Budil, P., Kraft, P., Fatka, O., and Marek, J., 2003, Provenance and early genetic processes of the Ordovi-
Nd-Sr-Pb isotopic signatures of Neoproterozoic–Early Paleozoic siliciclastic rocks cian Šárka Formation at Praha - Cˇervený vrch Hill (Barrandian, Czech Republic): Bulletin of Geosciences, v. 78, p. 147–156. Drost, K., Linnemann, U., McNaughton, N., Fatka, O., Kraft, P., Gehmlich, M., Tonk, C., and Marek, J., 2004, New data on the Neoproterozoic-Cambrian geotectonic setting of the Teplá-Barrandian volcano-sedimentary successions: Geochemistry, U-Pb zircon ages, and provenance (Bohemian Massif, Czech Republic): International Journal of Earth Sciences, v. 93, p. 742–757, doi: 10.1007/s00531-004-0416-5. Fatka, O., and Gabriel, Z., 1991, Microfossils from siliceous stromatolitic rocks of the Barrandian Proterozoic (Bohemian Massif): Cˇasopis pro mineralogii a geologii, v. 36, p. 143–148. Fatka, O., Frýda, J., Kachlík, V., Kraft, P., and Dolejš, D., 1998, Sedimentary and volcanic events in the Teplá-Barrandian crustal segment (Bohemia) as a response to dated intracrustal processes: Schriftenreihe des Staatlichen Museums für Mineralogie und Geologie zu Dresden, v. 9, p. 136–137. Fiala, F., 1971, Ordovician diabase volcanism and biotite lamprophyres of the Barrandian: Sborník Geologických Veˇd – Geologie, v. 19, p. 7–97 (English summary). Fiala, F., 1977, The Upper Proterozoic volcanism of the Barrandian area and the problem of spilites: Sborník Geologických Veˇd—Geologie, v. 30, p. 1–247. Fiala, F., 1978, The TiO2-K2O-P2O5 diagram and tectonomagmatic relations of the volcanics of the Barrandian area: Veˇstník ústrˇedního ústavu geologického, v. 53, p. 333–346. Friedl, G., Finger, F., Paquette, J.L., von Quadt, A., McNaughton, N.J., and Fletcher, I.R., 2004, Pre-Variscan geological events in the Austrian part of the Bohemian Massif deduced from U–Pb zircon ages: International Journal of Earth Sciences, v. 93, p. 802–823, doi: 10.1007/s00531-004-0420-9. Gehmlich, M., and Drost, K., 2005, Cadomian and Early Paleozoic magmatic events: Insights from zircon morphology and geochemical signatures of Iand S-type granitoids (Saxo-Thuringia/Bohemian Massif/central Europe): International Geology Review, v. 47, p. 1287–1297. Gerstenberger, H., and Haase, G., 1997, A highly effective emitter substance for mass spectrometric Pb isotope ratio determinations: Chemical Geology, v. 136, p. 309–312, doi: 10.1016/S0009-2541(96)00033-2. Gradstein, F.M., Ogg, J.G., Smith, A.G., eds., 2005, Geologic time scale 2004: Cambridge, Cambridge University Press, 610 p. Havlícˇek, V., 1971, Stratigraphy of the Cambrian of central Bohemia: Sborník Geologických Veˇd—Geologie, v. 20, p. 7–50. Havlícˇek, V., 1998, V. Prague basin—Ordovician, in Chlupácˇ, I., Havlícˇek, V., Krˇiž, J., Kukal, Z., and Štorch, P., eds., Palaeozoic of the Barrandian (Cambrian to Devonian): Prague, Czech Geological Survey, p. 41–79. Havlícˇek, V., and Fatka, O., 1992, Ordovician of the Prague basin (Barrandian area, Czechoslovakia), in Webby, B.D., and Laurie, J.R., eds., Global perspectives on Ordovician geology: Rotterdam, Balkema, p. 461–472. Havlícˇek, V., and Vaneˇk, J., 1966, The biostratigraphy of the Ordovician of Bohemia: Sborník Geologických Veˇd—Paleontologie, v. 8, p. 7–69. Holubec, J., 1995, VII.B.2 Structure, in Dallmeyer, D.R., Franke, W., and Weber, K., eds., Pre-Permian geology of central and eastern Europe: Berlin, Heidelberg, Springer, p. 392–397. Inglis, J.D., Samson, S.D., D’Lemos, R.S., and Hamilton, M., 2004, U-Pb geochronological constraints on the tectonothermal evolution of the Paleoproterozoic basement of Cadomia, La Hague, NW France: Precambrian Research, v. 134, p. 293–315, doi: 10.1016/j.precamres.2004.07.003. Jaffey, A.H., Flynn, K.F., Glendenin, L.E., Bentley, W.C., and Essling, A.M., 1971, Precision measurement of half-lives and specific activities of 235U and 238U: Physical Review C, v. 4, p. 1889–1906, doi: 10.1103/PhysRevC.4.1889. Jakeš, P., Zoubek, J., Zoubková, J., and Franke, W., 1979, Graywackes and metagraywackes of the Teplá-Barrandian Proterozoic area: Sborník Geologických Veˇd—Geologie, v. 33, p. 83–122. Jonas, P., Buschmann, B., and Gaitzsch, B., 2000, Unterkambrischer und unterkarboner Vulkanismus der Torgau-Doberlug Synklinale (NE Saxothuringische Zone): Zeitschrift für geologische Wissenschaften, v. 28, p. 157–175. Kettner, R., 1918, Versuch einer stratigraphischen Einteilung des böhmischen Algonkiums: Geologische Rundschau, v. 8, p. 169–188, doi: 10.1007/ BF01800895. Konzalová, M., 1981, Some late Precambrian microfossils from the Bohemian Massif and their correlation: Precambrian Research, v. 15, p. 43–62, doi: 10.1016/0301-9268(81)90066-8. Kraft, J., Kraft, P. and Fatka, O., eds., 1999, Excursion guide Barrandian: 8th International Symposium on the Ordovician System, Prague, 27 June– 1 July 1999, 60 p. Kraft, P., Linnemann, U., and Mazur, S., eds., 2004, Gondwanan-margin of the Rheic Ocean in the Bohemian Massif: Excursion guide books and abstracts, Opening meeting 2004 of the International Geological Correlation Program Project 497 “The Rheic Ocean: Its origin, evolution and
207
correlatives,” Prague, 20 September: Prague, Faculty of Science, Charles University Prague, 108 p. Krˇíbek, B., Pouba, Z., Skocˇek, V., and Waldhausrová, J., 2000, Neoproterozoic of the Teplá-Barrandian unit as a part of the Cadomian orogenic belt: A review and correlation aspects: Bulletin of the Czech Geological Survey, v. 75, p. 175–196. Krˇiž, J., 1998, V. Prague Basin—Silurian, in Chlupácˇ, I., Havlícˇek, V., Krˇiž, J., Kukal, Z., and Štorch, P., eds., Palaeozoic of the Barrandian (Cambrian to Devonian): Prague, Czech Geological Survey, p. 79–101. Kukal, Z., 1971, Sedimentology of Cambrian deposits of the Barrandian area (central Bohemia): Sborník Geologických Veˇd—Geologie, v. 20, p. 53–100. Lang, M., 2000, Composition of Proterozoic greywackes in the Barrandian. Veˇstník Cˇeského geologického ústavu, v. 75, p. 205–216. Liew, T.C., and Hofmann, A.W., 1988, Precambrian crustal components, plutonic associations, plate environment of the Hercynian fold belt of central Europe: Indications from a Nd and Sr isotopic study: Contributions to Mineralogy and Petrology, v. 98, p. 129–138, doi: 10.1007/BF00402106. Linnemann, U., and Romer, R.L., 2002, The Cadomian orogeny in SaxoThuringia, Germany: Geochemical and Nd-Sr-Pb isotopic characterization of marginal basins with constraints to geotectonic setting and provenance: Tectonophysics, v. 352, p. 33–64, doi: 10.1016/S0040-1951(02)00188-9. Linnemann, U., McNaughton, N.J., Romer, R.L., Gehmlich, M., Drost, K., and Tonk, C., 2004, West African provenance for Saxo-Thuringia (Bohemian Massif): Did Armorica ever leave pre-Pangean Gondwana?—U/PbSHRIMP zircon evidence and the Nd-isotopic record: International Journal of Earth Sciences, v. 93, p. 683–705, doi: 10.1007/s00531-004-0413-8. Lüning, S., Craig, J., Loydell, D.K., Štorch, P., and Fitches, B., 2000, Lower Silurian “hot shales” in north Africa and Arabia: Regional distribution and depositional model: Earth Science Reviews, v. 49, p. 121–200, doi: 10.1016/S0012-8252(99)00060-4. Manhès, G., Minster, J.F., and Allègre, C.J., 1978, Comparative uranium-thorium-lead and rubidium-strontium study of the Saint Séverin amphoterite: Consequences for early solar system chronology: Earth and Planetary Science Letters, v. 39, p. 14–24, doi: 10.1016/0012-821X(78)90137-1. Mašek, J., 2000, Stratigraphy of the Proterozoic of the Barrandian area: Veˇstník Cˇeského geologického ústavu, v. 75, p. 197–200. McCaffrey, W.D., 1994, Sm-Nd isotopic characteristics of sedimentary provenance: The Windermere Supergroup of NW England: Journal of the Geological Society of London, v. 151, p. 1017–1021. Murphy, J.B., and Nance, R.D., 1989, Model for the evolution of the Avalonian-Cadomian belt: Geology, v. 17, p. 735–738, doi: 10.1130/00917613(1989)017<0735:MFTEOT>2.3.CO;2. Murphy, J.B., and Nance, R.D., 2002, Sm-Nd isotopic systematics as tectonic tracers: An example from West Avalonia in the Canadian Appalachians: Earth Science Reviews, v. 59, p. 77–100, doi: 10.1016/S00128252(02)00070-3. Murphy, J.B., Pisarevsky, S.A., Nance, R.D., and Keppie, J.D., 2004, Neoproterozoic–Early Paleozoic evolution of peri-Gondwanan terranes: Implications for Laurentia-Gondwana connections: International Journal of Earth Sciences, v. 93, p. 659–682, doi: 10.1007/s00531-004-0412-9. Nägler, T.F., Schäfer, H.J., and Gebauer, D., 1995, Evolution of the western European continental crust: Implications from Nd and Pb isotopes in Iberian sediments: Chemical Geology, v. 121, p. 345–357, doi: 10.1016/00092541(94)00129-V. Nance, R.D., and Murphy, J.B., 1994, Contrasting basement isotopic signatures and the palinspastic restoration of peripheral orogens: Example from the Neoproterozoic Avalonian-Cadomian belt: Geology, v. 22, p. 617–620, doi: 10.1130/0091-7613(1994)022<0617:CBISAT>2.3.CO;2. Nance, R.D., and Murphy, J.B., 1996, Basement isotopic signatures and Neoproterozoic paleogeography of Avalonian-Cadomian and related terranes in the circum North Atlantic, in Nance, R.D., and Thompson, M.D., eds., Avalonian and related peri-Gondwanan terranes of the circum-North Atlantic: Boulder, Colorado, Geological Society of America Special Paper 304, p. 333–346. Nance, R.D., Murphy, J.B., Strachan, R.A., D’Lemos, R.S., and Taylor, G.K., 1991, Late Proterozoic tectonostratigraphic evolution of the Avalonian and Cadomian terranes: Precambrian Research, v. 53, p. 41–78, doi: 10.1016/0301-9268(91)90005-U. Nesbitt, H.W., and Young, G.M., 1982, Early Proterozoic climates and plate motions inferred from major element chemistry of lutites: Nature, v. 299, p. 715–717, doi: 10.1038/299715a0. Nesbitt, H.W., Markovics, G., and Price, R.C., 1980, Chemical processes affecting alkalis and alkaline earths during continental weathering: Geochimica et Cosmochimica Acta, v. 44, p. 1659–1666, doi: 10.1016/00167037(80)90218-5.
208
Drost et al.
Neuroth, H., 1997, K/Ar-Datierungen an detritschen Muskoviten—“Sicherungskopien” orogener Prozesse am Beispiel der Varisziden: Göttinger Arbeiten zur Geologie und Paläontologie, v. 72, p. 1–143. Pacltová, B., 1990, Late Proterozoic organic remains from the Mítov locality, in Chanda, S., ed., Silver Jubilee commemoration: Journal of Palynology, v. 91, p. 261–276. Pašava, J., and Amov, B., 1993, Isotopic composition of lead in Proterozoic anoxic metasedimentary and volcanogenic rocks from the Bohemian Massif (Czech Republic) with metallogenetic implications: Chemical Geology, v. 109, p. 293–304, doi: 10.1016/0009-2541(93)90076-U. Patocˇka, F., and Štorch, P., 2004, Evolution of geochemistry and depositional settings of Early Palaeozoic siliciclastics of the Barrandian (Teplá-Barrandian unit, Bohemian Massif, Czech Republic): International Journal of Earth Sciences, v. 93, p. 728–741. Patocˇka, F., Vlašímský, P., and Blechová, K., 1993, Geochemistry of early Paleozoic volcanics of the Barrandian basin (Bohemian Massif, Czech Republic): Implications for paleotectonic reconstructions: Jahrbuch der Geologischen Bundesanstalt, v. 136, p. 873–896. Pelc, Z., and Waldhausrová, J., 1994, Geochemical characteristics of volcanics of the Barrandian Upper Proterozoic in the Chudenice-Nepomuk region (SW Bohemia): Sborník Geologických Veˇd—Geologie, v. 46, p. 5–21. Pin, C., and Waldhausrová, J., 2007 (this volume), Sm-Nd isotope and trace element study of the 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: Boulder, Colorado, Geological Society of America Special Paper 423, doi: 10.1130/2007.2423(10). Pin, C., Kryza, R., Oberc-Dziedzic, T., Mazur, S., Turniak, K., and Waldhausrová, J., 2007 (this volume), 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, 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: Boulder, Colorado, Geological Society of America Special Paper 423, doi: 10.1130/2007.2423(09). Pupin, J.P., 1980, Zircon and granite petrology: Contributions to Mineralogy and Petrology, v. 73, p. 207–220, doi: 10.1007/BF00381441. Röhlich, P., 1965, Geologische Probleme des mittelböhmischen Algonkiums: Geologie, v. 14, p. 373–403. Röhlich, P., 2000, Some stratigraphic problems of the Barrandian Neoproterozoic: Veˇstník Cˇeského geologického ústavu, v. 75, p. 201–204. Romer, R.L., Förster, H.-J., and Breitkreuz, C., 2001, Intracontinental extensional magmatism with a subduction fingerprint: The late Carboniferous Halle volcanic complex (Germany): Contributions to Mineralogy and Petrology, v. 141, p. 201–221. Romer, R.L., Heinrich, W., Schröder-Smeibidl, B., Meixner, A., Fischer, C.-O., and Schulz, C., 2005, Elemental dispersion and stable isotope fractionation during reactive fluid-flow and fluid immiscibility in the Bufa del Diente aureole, NE-Mexico: Evidence from radiographies and Li, B, Sr, Nd, and Pb isotope systematics: Contributions to Mineralogy and Petrology, v. 149, p. 400–429, doi: 10.1007/s00410-005-0656-9. Roser, B.P., and Korsch, R.J., 1986, Determination of tectonic setting of sandstone-mudstone suites using SiO2 content and K2O/Na2O ratio: Journal of Geology, v. 94, p. 635–650. Samson, S.D., and D’Lemos, R.S., 1998, U-Pb geochronology and Sm-Nd isotopic composition of Proterozoic gneisses, Channel Islands, UK: Journal of the Geological Society of London, v. 155, p. 609–618. Sánchez-García, T., Bellido, F., and Quesada, C., 2003, Geodynamic setting and geochemical signatures of Cambrian-Ordovician rift-related igneous rocks (Ossa-Morena Zone, SW Iberia): Tectonophysics, v. 365, p. 233– 255, doi: 10.1016/S0040-1951(03)00024-6. Simien, F., Mattauer, M., and Allègre, C.J., 1999, Nd isotopes in the stratigraphic record of the Montagne Noire (French Massif Central): No significant Paleozoic juvenile inputs, and pre-Hercynian paleogeography: Journal of Geology, v. 107, p. 87–97, doi: 10.1086/314332. Sláma, J., Kachlík, V., and Košler, J., 2003, Neoproterozoic metaconglomerates of the Sedlcˇany-Krásná Hora Metamorphic Islet: GeoLines, v. 16, p. 96–97. Steiger, R.H., and Jäger, E., 1977, Subcommission on geochronology: Convention on the use of decay constants in geo- and cosmochronology: Earth and Planetary Science Letters, v. 36, p. 359–362, doi: 10.1016/0012821X(77)90060-7. Štorch, P., 1986, Ordovician-Silurian boundary in the Prague basin (Barrandian area, Bohemia): Sborník Geologických Veˇd—Geologie, v. 41, p. 69–103.
Štorch, P., 1990, Upper Ordovician–Lower Silurian sequences of the Bohemian Massif, central Europe: Geological Magazine, v. 127, p. 225–239. Štorch, P., 1998, VIII. Volcanism, in Chlupácˇ, I., Havlícˇek, V., Krˇiž, J., Kukal, Z., and Štorch, P., eds., Palaeozoic of the Barrandian (Cambrian to Devonian): Prague, Czech Geological Survey, p. 149–164. Štorch, P., Fatka, O., and Kraft, P., 1993, Lower Paleozoic of the Barrandian area (Czech Republic)—A review: Coloquios de Paleontología, Madrid, v. 45, p. 163–191. Strnad, L., and Mihaljevicˇ, M., 2005, Sedimentary provenance of Mid-Devonian clastic sediments in the Teplá-Barrandian unit (Bohemian Massif): U-Pb and Pb-Pb geochronology of detrital zircons by laser ablation ICP-MS: Mineralogy and Petrology, v. 84, p. 47–68, doi: 10.1007/s00710-004-0057-1. Taylor, S.R., and McLennan, S.M., 1985, The continental crust: Its composition and evolution: Blackwell Scientific Publications, 312 p. Thorogood, E.J., 1990, Provenance of the pre-Devonian sediments of England and Wales: Sm-Nd isotopic evidence: Journal of the Geological Society of London, v. 147, p. 591–594. Tilton, G.R., 1973, Isotopic lead ages of chondritic meteorites: Earth and Planetary Science Letters, v. 19, p. 321–329, doi: 10.1016/0012-821X(73)90082-4. Vejnar, Z., 1966, The petrogenetic interpretation of kyanite, silimanite and andalusite in the southwestern Bohemian crystalline complexes: Neues Jahrbuch für Mineralogie Abhandlungen, v. 104, p. 172–189. Venera, Z., Schulmann, K., and Kröner, A., 2000, Intrusion within a transitional tectonic domain: The Cˇistá granodiorite (Bohemian Massif)—Structure and rheological modelling: Journal of Structural Geology, v. 22, p. 1437– 1454, doi: 10.1016/S0191-8141(00)00054-7. Vidal, P., Auvray, B., Charlot, R., Fediuk, F., Hameurt, J., and Waldhausrová, J., 1975, Radiometric age of volcanics of the Cambrian “KrˇivoklátRokycany” complex (Bohemian Massif): Geologische Rundschau, v. 64, p. 563–570, doi: 10.1007/BF01820683. Vokurka, K., and Frýda, J., 1997, The neodymium isotopes in Lower Paleozoic basalts from the Barrandian (Teplá-Barrandian unit, Bohemian Massif): Zprávy o geologických výzkumech v roce 1996, p. 87. Waldhausrová, J., 1971, The chemistry of the Cambrian volcanics in the Barrandian area: Krystalinikum, v. 8, p. 45–75. Waldhausrová, J., 1984, Proterozoic volcanics and intrusive rocks of the Jílové zone in central Bohemia: Krystalinikum, v. 17, p. 77–97. Waldhausrová, J., 1997a, Geochemistry of volcanics (metavolcanics) in the western part of the TBU Precambrian and their original geotectonic setting, in Vrána, S., and Šteˇdrá, V., eds., Geological model of western Bohemia related to the KTB borehole in Germany: Journal of Geological Sciences, v. 47, p. 85–90. Waldhausrová, J., 1997b, Proterozoic volcanics geochemistry and mineral chemistry: A contribution to the Barrandian upper Proterozoic stratigraphy (Bohemian Massif, Czech Republic): Krystalinikum, v. 23, p. 151–180. Wendt, J.I., Kröner, A., Fiala, J., and Todt, W., 1993, Evidence from zircon dating for existence of approximately 2.1 Ga old crystalline basement in southern Bohemia, Czech Republic: Geologische Rundschau, v. 82, p. 42–50, doi: 10.1007/BF00563269. Wronkiewicz, D.J., and Condie, K.C., 1987, Geochemistry of Archean shales from the Witwatersrand Supergroup, South Africa: Source-area weathering and provenance: Geochimica et Cosmochimica Acta, v. 51, p. 2401– 2416, doi: 10.1016/0016-7037(87)90293-6. Zulauf, G., 1997, Von der Anchizone bis zur Eklogitfazies: Angekippte Krustenprofile als Folge der cadomischen und variscischen Orogenese im Teplá-Barrandium (Böhmische Masse): Geotektonische Forschungen, v. 89, p. 302. Zulauf, G., Dörr, W., Fiala, J., and Vejnar, Z., 1997a, Lower Carboniferous lift tectonics in the Bohemian Massif: A special kind of gravitational collapse in thickened orogenic crust: Terra Nostra, v. 97, no. 5, p. 216–218. Zulauf, G., Dörr, W., Fiala, J., and Vejnar, Z., 1997b, Late Cadomian crustal tilting and Cambrian transtension in the Teplá-Barrandian unit (Bohemian Massif, central European Variscides): Geologische Rundschau, v. 86, p. 571–587, doi: 10.1007/s005310050164. 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. 150/4, p. 627–639. Zulauf, G., Bues, C., Dörr, W., and Vejnar, Z., 2002, 10 km minimum throw along the West Bohemian shear zone: Evidence for dramatic crustal thickening and high topography in the Bohemian Massif (European Variscides): International Journal of Earth Sciences, v. 91, p. 850–864, doi: 10.1007/s00531-001-0250-y. MANUSCRIPT ACCEPTED BY THE SOCIETY 3 OCTOBER 2006 Printed in the USA
Geological Society of America Special Paper 423 2007
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 Christian Pin* Département de Géologie, UMR 6524 CNRS, Université Blaise Pascal, 5 rue Kessler, 63 038 Clermont-Ferrand, France R. Kryza T. Oberc-Dziedzic S. Mazur K. Turniak Institute of Geological Sciences, University of Wrocław, plac Maxa Borna 9, 50-204 Wrocław, Poland Jarmila Waldhausrová Nad Hercovkow 422, 18200 Prague, Czech Republic
ABSTRACT Ca. 500 Ma orthogneisses and bimodal suites are widespread along the northern part of the Bohemian Massif (central European Variscides) and are interpreted to document intense magmatism during a continental break-up episode along the northern periphery of Gondwana. Based on geological setting, and geochemical and isotopic evidence, these felsic igneous rocks record the generation of: (1) magmas of pure or predominantly crustal derivation, represented by minor extrusives and much more voluminous orthogneisses similar to S-type granitoids; (2) subordinate magmas of exclusively mantle origin (ranging from within-plate alkali trachytes to oceanic plagiogranites) corresponding to felsic derivatives of associated basalts; and (3) magmas of hybrid origin, produced either as a result of large degrees of contamination of mantle-derived magmas ascending through the crust, or alternatively, generated by partial melting of mixed sources, such as interlayered sediments and mafic rocks or graywackes containing a juvenile component. The high-temperature dehydration melting process responsible for the generation of the most abundant rock-types necessitated the advection of mantle heat, in a context of continental lithosphere extension, as documented by broadly coeval basaltic magmatism at the scale of the igneous province. The large volumes of felsic magmas generated during the 500-Ma anorogenic event are interpreted to result from the combination of a hot extensional tectonic regime with the widespread availability in the lower crust of fertile lithologies, such as metagraywackes. This in turn reflects the largely undifferentiated nature of the crustal segment accreted some 50–100 m.y. earlier during the Cadomian orogeny. Keywords: granitoids, A-type, geochemistry, partial melting, fertility *E-mail:
[email protected]. Pin, C., Kryza, R., Oberc-Dziedzic, T., Mazur, S., Turniak, K., and Waldhausrová, J., 2007, 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, 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. 209–229, doi: 10.1130/2007.2423(09). For permission to copy, contact
[email protected]. ©2007 Geological Society of America. All rights reserved.
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INTRODUCTION
Orthogneisses
Geochronological data, especially precise U-Pb zircon ages, gathered throughout the Bohemian Massif during the past 15 years have documented the existence of a widespread igneous episode of Late Cambrian to Early Ordovician age (ca. 510–490 Ma). Besides mafic and bimodal associations, copious volumes of orthogneisses document the production of large amounts of felsic magmas, the origin of which has proved to be controversial. Based on alleged broadly “calcalkaline” geochemical features and the use of some trace element discrimination diagrams, some authors (e.g., Oliver et al., 1993; Kröner and Hegner, 1998, Kröner et al., 2001) interpreted these meta-granitoids as remnants of a batholith emplaced above a subduction zone in a continental active margin setting, related to the closure of the northern (in presentday coordinates) Tornquist Ocean separating Baltica from a supposed eastern prolongation of Avalonia. In contrast, other authors (e.g., Kryza and Pin, 1997; Crowley et al., 2000, 2002; Floyd et al., 2000; Dostal et al., 2001) emphasized the absence of intermediate (andesitic) rock-types typical of orogenic magmatism. These authors used the broad spatial and temporal association of within-plate mafic rocks to suggest that granite magmatism occurred in rift-related geodynamic environments, in a large-scale context of continental break-up leading to the opening of an ocean basin (Rheic) within the so-called “Armorica realm,” at the northern edge of Gondwana. In this article, we present new Sm-Nd isotope results obtained from ca. 500-Ma metagranitoids from the Polish Sudetes and from coeval volcanics of the Barrandian area, and review geochemical and isotope data available on broadly contemporaneous igneous rocks from the northern part of the Bohemian Massif. These data highlight the diversity of felsic magmas produced during the ca. 500-Ma event, and can be used to put constraints on possible source materials and draw inferences on the petrogenetic processes responsible for magma generation and their tectonic setting.
Orlica-S´niez˙nik Massif The Orlica-S´niez˙nik Massif comprises predominantly amphibolite-grade orthogneisses and staurolite-grade mica schists, and contains inclusions of high-pressure and ultra-high-pressure rocks. Supracrustal series of presumably Neoproterozoic to Cambrian age (e.g., Don et al., 1990) were intruded by large granitic plutons at ca. 500 Ma (Oliver et al., 1993; Turniak et al., 2000; Kröner et al., 2001). The subsequent Variscan tectonothermal processes were complex and included medium- to high-grade metamorphism and intense synmetamorphic deformation, accompanied by exhumation of high-grade rocks (e.g., Bröcker and Klemd 1996; Lange et al., 2002, 2005; Štipská et al., 2004, and references therein). The Orlica-S´niez˙nik Massif has usually been considered to represent a gneissic dome, in which the gneisses crop out in antiforms, whereas “mantling” schists are preserved in synforms. There is some evidence, however, that the Orlica-S´niez˙nik Massif is composed of a number of folded thrust sheets, as are the East Sudetic units forming a nappe pile adjacent to the southeast. The ca. 500-Ma granitoids were transformed into orthogneisses known as two main varieties: the S´niez˙nik and Gierałtów gneisses (e.g., Borkowska et al., 1990; Don et al., 1990, and references therein). Typically, the S´niez˙nik gneisses are coarse- to medium-grained rocks with K-feldspar megacrysts, composed of quartz, microcline, oligoclase, muscovite, biotite and accessory apatite, zircon, rutile, titanite (late-stage), garnet, and opaque minerals. Local transitions to well-layered textural varieties are interpreted to reflect higher strain intensity (Turniak et al., 2000; Lange et al., 2002). In contrast, the Gierałtów gneisses are finegrained, approximately equigranular, often thinly laminated twomica rocks of granite composition similar to that of the S´niez˙nik gneisses. Locally, they display a migmatitic texture, considerably different from the augen texture of the S´niez˙nik gneisses. The Gierałtów gneisses are accompanied by lensoid bodies of amphibolites, granulites, and eclogites. Transitional textural varieties occur between the S´niez˙nik and Gierałtów gneisses. Sensitive high-resolution ion microprobe (SHRIMP) analyses of zircons from both kinds of gneisses yielded ages of ca. 500 Ma, interpreted to reflect igneous crystallization from similar magmas, as also suggested by common mineralogical and geochemical characteristics (Turniak et al., 2000). Specifically, both gneiss varieties are peraluminous in character, as indicated by their mineralogy (two micas, minor garnet, and rare cordierite), the Al2O3/ (CaO + Na2O + K2O) molecular ratios (A/CNK) ranging from 1.01 to 1.22, and high normative corundum, between 1.15 and 2.73. Their trace element patterns are broadly similar, with strong light rare earth element (LREE) fractionation and flat heavy rare earth element (HREE) distribution (Fig. 2); all samples display distinct negative anomalies for Nb, Sr, and Ti (Fig. 2). Both the orthogneisses were interpreted to have derived from geochemically similar protoliths (Lange et al., 2002, 2005) from mature crust, but are difficult to constrain in terms of geodynamic environments (Turniak et al., 2000).
GEOLOGICAL CONTEXT AND REVIEW OF GEOCHEMICAL AND Sm-Nd DATA The location of the ca. 500-Ma igneous suites referred to in this study is shown in the geological framework of the northern part of the Bohemian Massif (Fig. 1). The geological context and major feature of the various occurrences of metagranitoids and metavolcanics are briefly described from the east (Sudetes) to the west (Saxo-Thuringian zone), based on published studies and some new Sm-Nd data (Orlica-S´niez˙nik Massif gneisses; Leszczyniec metavolcanics). A short account is then given of the geological setting of the nonmetamorphic volcanics of the Krˇ ivoklát-Rokycany Complex (Barrandian), together with an assessment of published and new geochemical and isotope data.
Late Cambrian (ca. 500 Ma) felsic anorogenic magmatism in the Bohemian Massif se
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Figure 1. Geological sketch of the northern part of the Bohemian Massif. The areas studied and described in this article are shaded (500-Ma orthogneisses and volcanic complexes). EK—East Karkonosze; FG—Fichtelgebirge; FSB—fore-Sudetic block; GSM—Góry Sowie Massif; ISF—intra-Sudetic fault zone; KC—Kaczawa complex; KG—Karkonosze Granite; KIM—Karkonosze-Izera Massif; KRC—Krˇivoklát-Rokycany complex; OSM—Orlica-S´niez˙nik Massif; SK—south Karkonosze; VC—Vesser complex.
New Sm-Nd data obtained on a subset of the samples studied by Turniak et al. (2000) are listed in Table 1. All the samples yield clearly negative initial εNd values (from −3.5 to −6.4), demonstrating that their source material was depleted in Sm relative to Nd on a time-integrated basis. This depletion suggests that ancient crustal materials enriched in LREE played a major role in the genesis of these rocks, as suggested by crustal residence model ages. These scatter from 1.4 to 2.1 Ga, far in excess of the emplacement age (samples MS14 and MS17, whose very high 147 Sm/144Nd ratios probably reflect late-stage fractionation of a LREE-rich phase, such as monazite, were excluded from TDM calculations). Although a broad homogeneity is indicated, it can be noticed that the three S´niez˙nik samples tend to have slightly higher εNd values than those for the Gierałtów gneisses, suggesting that the augen gneisses might have been extracted from slightly less evolved crustal sources than their equigranular counterparts. Our
Nd isotope data are in general agreement with the results published by Kröner et al. (2001) for six gneiss samples from the Orlica-S´niez˙nik Massif and confirm the clear crustal derivation of the ca. 500-Ma granitoids. We do not concur, however, with their interpretation in terms of “extensive melting of Precambrian basement,” nor with their suppositions on the subduction-related nature of these granitoids (see Discussion section). Góry Sowie Massif The Góry Sowie Massif in the central Sudetes is mostly composed of gneisses and migmatites with minor intercalations of amphibolites, as well as small bodies of high-temperature–highpressure (HT-HP) felsic granulites and associated meta-ultramafic rocks (for an overview, see Kryza, 1981; Z˙ elaz´niewicz, 1990). The gneisses and migmatites comprise a wide range of textural varieties and typically contain an assemblage of quartz
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Figure 2. Selected trace element diagrams for representative samples of ca. 500-Ma orthogneisses from the northern part of the Bohemian Massif (normalized to chondrites, after Nakamura, 1974, and Thompson, 1982). Diagram made using the GCDkit software of Janoušek et al. (2003). Data sources: Góry Sowie Massif—Kröner and Hegner (1998); Orlica-S´niez˙nik Massif—Turniak et al. (2000); Izera Massif—Oberc-Dziedzic et al. (2005a); Strzelin-Lipowe massifs—Oberc-Dziedzic et al. (2005b); Erzgebirge—Tichomirowa et al. (2001), Mingram et al. (2004); Fichtelgebirge—Siebel et al. (1997). FSB—fore-Sudetic block. The symbols of the samples are from the original papers cited.
Late Cambrian (ca. 500 Ma) felsic anorogenic magmatism in the Bohemian Massif ´ NIEZ˙NIK TABLE 1. Sm-Nd ISOTOPE DATA FOR THE ORLICA-S MASSIF ORTHOGNEISSES 147 144 143 144 Sample Orthogneiss Sm Nd Sm/ Nd Nd/ Nd 2 εNd500Ma type S.E. ´ niez˙nik MS5 S 7.32 35.9 0.1234 0.512220 7 –3.5 ´ niez˙nik MS10 S 3.13 13.7 0.1380 0.512166 7 –5.5 ´ niez˙nik MS20 S 6.69 31.2 0.1298 0.512187 8 –4.6 ´ niez˙nik? OB1 S 5.47 24.0 0.1379 0.512163 7 –5.6 L1 S´niez˙nik? 5.98 26.4 0.1369 0.512193 7 –4.9 MS9 Gierałtów 6.15 28.3 0.1315 0.512185 6 –4.7 MS14 Gierałtów 1.76 6.14 0.1737 0.512261 10 –5.9 MS16 Gierałtów 4.29 17.4 0.1492 0.512159 15 –6.4 MS17 Gierałtów 2.83 9.03 0.1896 0.512298 9 –6.2 MS18 Gierałtów 4.59 21.6 0.1285 0.512192 7 –4.4 Note: S.E.—standard error; – indicates data not available.
+ plagioclase + biotite ± muscovite ± garnet ± sillimanite. Most of the gneisses and migmatites were usually interpreted as derived from sedimentary protoliths, based on field and petrological constraints (e.g., lithological variation, widespread and often abundant sillimanite). However, relatively limited rock varieties (Fig. 1), two-feldspar gneisses, are petrographically rather homogeneous and often display augen texture, and they may have derived from granites. The geochronology of the Góry Sowie Massif rocks caused much controversy. However, common varieties of the gneisses and migmatites (layered and diatexitic types) were described as containing variable zircon populations, but with one dominant group of magmatic-type zircons (prismatic and oscillatory zoned) dated at ca. 500 Ma (Kryza and Fanning, 2007). These magmatic-type zircons were controversially interpreted either as evidence for igneous protoliths of the gneisses (Kröner and Hegner, 1998) or as representing both the igneous protoliths and sedimentary materials from reworked igneous sources (Kryza and Fanning, 2007). Representative samples of the Góry Sowie Massif gneisses and migmatites were studied for major and trace elements and Sm-Nd isotopes by Kröner and Hegner (1998), who interpreted the vast majority of the rocks as orthogneisses derived from magmatic arc–type granitoids. However, detailed geochemical and isotopic studies on the associated mafic rocks (amphibolites, coronitic metagabbros) revealed a predominance of enriched mantle-derived magmas emplaced in an extensional (probably immature rift) setting (Winchester et al., 1998; Kryza and Pin, 2002), also suggesting an extensional emplacement setting for the enclosing gneisses. The amphibolite-facies metamorphism and migmatization of the Góry Sowie Massif gneisses took place at ca. 385–370 Ma, as evidenced by isotopic ages obtained by various methods (see Kryza et al., 2004, and references therein). In the set of samples studied by Kröner and Hegner (1998), two specimens, PL16 and PL17, represent augen gneisses of the most likely magmatic derivation. Their modal composition (quartz, K-feldspar, plagioclase, muscovite, and biotite) and major element characteristics are typical of granites, whereas the trace element distributions display strong LREE and weak HREE
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TDM (Ma) 1390 1770 1555 1770 1690 1590 3085 2095 – 1525
fractionations, with distinct negative Nb, Sr, Ti, and Eu anomalies (Fig. 2). Two other samples of layered migmatites (PL4 and PL6) and one representing diatexite (PL11) differ from the augen gneisses in having no K-feldspar in their mode but containing accessory garnet and sillimanite; their trace element patterns are roughly similar to those of the augen gneisses. In all the samples, magmatic-type zircon populations (euhedral prismatic habit and strong oscillatory zonation) of ca. 500-Ma age are common (Kröner and Hegner, 1998; Kryza and Fanning, 2007). The Góry Sowie augen gneisses (PL16 and PL17) show εNd values of +0.2 and TDM of 1.13, whereas the migmatites display negative εNd values between –3.0 and –6.2, and TDM up to 1.56 (Kröner and Hegner, 1998). Three other samples from the Góry Sowie rocks were also studied for Sm-Nd isotopes by Crowley et al. (2002), giving broadly similar characteristics (unfortunately petrographic features and locations of the samples were not given). Izera-Karkonosze Massif The Izera-Karkonosze Massif in the west Sudetes is composed of the Karkonosze Granite pluton (Fig. 1), which was emplaced 328 ± 12 Ma (Rb-Sr whole-rock isochron; Pin et al., 1987), and its metamorphic envelope, which is divided by the granite into two contrasting parts: the northern part, represented by the Izera Block, and the southeastern part, which comprises metamorphic complexes of the east and south Karkonosze. The Izera Block is composed mainly of porphyritic coarsegrained Izera granites and augen gneisses. Fine-grained granites and gneisses are minor throughout the area, whereas granodioritic gneisses are typical of the western part of the Izera Block. All textural types of the gneisses enclose lenses of coarse-grained porphyritic granites, mostly considered as remnants of undeformed granitic protoliths. The gneisses and granites locally contain small (up to several meters thick) younger but variously deformed basic dikes. The granitoid protoliths of the Izera gneisses gave igneous emplacement ages in the range of 515–480 Ma (Borkowska et al., 1980; Korytowski et al., 1993; Oliver et al., 1993; Z˙ elaz´niewicz, 1994; Kröner et al., 2001). The undeformed relics of these rocks correspond to the so-called “Rumburk granites” that intruded
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the 590–540-Ma (Z˙ elaz´niewicz et al., 2003) granodiorites of the Lusatian Massif on the west. The ca. 500-Ma granitoids of the Izera Block intruded metasedimentary and metavolcanic rocks (Berg, 1923; Achramowicz and Z˙ elaz´niewicz, 1998) of the Neoproterozoic age (640–550 Ma; Z˙ elaz´niewicz et al., 2003), which are preserved as three belts of mica schists, up to several hundred meters wide. The Izera Block suffered three deformation episodes (Mazur and Kryza, 1996; Z˙ elaz´niewicz et al., 2003), under mid- to low-grade metamorphic conditions in Late Devonian and Early Carboniferous times. In the eastern and southern Izera-Karkonosze Massif, the Kowary gneiss, dated at ca. 492–481 Ma (U-Pb ages; Oliver et al., 1993) and the Karkonosze gneiss (ca. 503 Ma; Kröner et al., 2001), both correspond to the Izera gneiss, whereas predominantly metasedimentary sequences (the so-called “Czarnów Formation” and the Velká Úpa Group) are equivalents of the mica schists of the Izera Block (Mazur and Aleksandrowski, 2001). The Izera granitoids are typical S-type granites, as documented by petrographic features (magmatic cordierite, relict garnet, sillimanite, the lack of mafic enclaves) and geochemistry: peraluminous character (A/CNK, 1.0–1.63), high normative corundum (up to 3.5%) and relatively high 87Sr/86Sr. Although they show some geochemical similarities to syn- and postcollisional granitoids, they were considered more likely to belong, based on a range of petrographic evidence, to the anorogenic class (Oberc-Dziedzic et al., 2005a). The εNd values for the Izera orthogneisses range from –5.2 to –6.9, implying that the magmas were extracted from a source with a secular enrichment in LREE. Their TDM ages, ranging between 1730 and 2175 Ma, point to an old crustal residence age of the inferred metasedimentary protoliths (Oberc-Dziedzic et al., 2005a). Fore-Sudetic Block In the fore-Sudetic block at the northeastern edge of the Bohemian Massif (FSB in Fig. 1), orthogneisses are known from vicinities of Strzelin and Wa˛droz˙e Wielkie, in the eastern and central parts of the block, respectively. In the former area, ca. 500Ma orthogneisses (Oliver et al., 1993; Kröner and Mazur, 2003; Oberc-Dziedzic et al., 2003b) are tectonically juxtaposed with older gneisses of ca. 600–568 Ma (Oberc-Dziedzic et al., 2003a). The younger gneisses were proved as derived from peraluminous S-type granitoids, and their trace element patterns are very similar to most of the orthogneisses described from the other areas (Oberc-Dziedzic et al., 2005b, and references therein; Fig. 2). Unfortunately, no Sm-Nd isotope data are yet available from these rocks. The orthogneisses from the other area of Wa˛droz˙e Wielkie, which are petrographically similar to the Izera gneisses, have recently been dated at ca. 540 Ma (Z˙ elaz´niewicz et al., 2004), and thus, as considerably older, are not considered here in detail. Erzgebirge (Eastern Saxo-Thuringian Zone) The Erzgebirge, in the eastern part of the Saxo-Thuringian zone (Fig. 1), exposes various types of gneisses, traditionally
grouped into red and gray varieties and interpreted as para- and orthogneisses, respectively. However, this subdivision and interpretation have appeared to be too simplified (Mingram et al., 2004, and references therein). Referring to recent structural and petrological studies (Kröner et al., 1995; Schmädicke and Evans, 1997; Willner et al., 1997; Rötzler et al., 1998; Tichomirowa et al., 2001; Mingram et al., 2004), the orthogneisses are found within the following three tectonometamorphic units (from bottom to top): (1) mediumpressure–medium-temperature (MP-MT) gneiss unit (granite gneisses), (2) HP-HT gneiss/eclogite unit (granulite facies orthogneisses), and (3) high-pressure–low-temperature (HP-LT) mica schist/eclogite unit (granite orthogneisses). Metagranitoid bodies occur also within the so-called “transition zones” between the units mentioned above and are interpreted as large-scale shear zones (Mingram et al., 2004, and references therein). Geochronological studies, combined with geochemical and Sm-Nd isotopic characteristics (Kröner et al., 1995; Tichomirowa et al., 2001; Mingram et al., 2004), indicate that at least two discrete magmatic events are contained in the red gneisses: (1) at ca. 550 Ma in MP-MT gneiss unit, and (2) at ca. 500–480 Ma in the high-pressure units. Based on mineralogy and textures, the Red Gneiss Group has been subdivided into three main rock types, namely, granite gneiss, augen gneiss, and muscovite gneiss (Mingram et al., 2004). The granite gneisses vary from coarse-grained porphyritic rocks to sheared mylonitic metagranitoids. They consist of K-feldspar, quartz, plagioclase, biotite, muscovite, and small amounts of garnet. Single zircons yielded Pb-Pb protolith ages between 550 and 560 Ma (Kröner et al., 1995). The augen gneisses contain large K-feldspar porphyrocrysts within a strained matrix of plagioclase, biotite, and muscovite. Samples of augen gneisses at the rim of Reitzenhain granite yielded mean Pb-Pb ages of 551 ± 9 (Kröner et al., 1995) and 492 ± 14 Ma (Tichomirowa, 2003, in Mingram et al., 2004). The muscovite gneisses occur in the HP units and the transition zones. They are mostly fine-grained rocks composed of quartz, K-feldspar, plagioclase, white mica, garnets, and biotite (± kyanite). Single zircons yielded a mean 207Pb/206Pb age of 479 ± 1 Ma (Kröner and Willner, 1998). Observed transitions between muscovite gneisses and metasedimentary rocks suggest that some of them could have derived from volcaniclastic deposits (Mingram et al., 2004). Geochemically, all granite gneisses and augen gneisses are broadly similar, corresponding to peraluminous granitoids, with moderate LREE fractionation (100–200 × chondrites) relative to HREE. The muscovite gneisses are significantly different, with higher K2O/Na2O of ~1.5, and lower Al2O3 contents (<14 wt%); their normalized REE diagrams display a spectrum of subparallel patterns with slight LREE enrichment (20–60 × chondrites) and flat HREE distribution (Mingram et al., 2004). Initial εNd values in all orthogneiss samples from the Erzgebirge range between −4.1 and −9.2, corresponding to crustal source with average residence times of 1.9–1.5 Ga (Mingram et al., 2004). These results lend support to the petrogenetic scenario proposed by Tichomirowa et
Late Cambrian (ca. 500 Ma) felsic anorogenic magmatism in the Bohemian Massif al. (2001), in which the ca. 500-Ma granitoids might have been generated through partial melting in the deep crust of metagraywackes and/or of ca. 540-Ma S-type granitoids. Fichtelgebirge (Western Saxo-Thuringian Zone) The Fichtelgebirge (northeast Bavaria) corresponds to a Late Variscan anticlinal zone in the northwestern part of the Bohemian Massif. It is mostly composed of low- to medium-grade detrital metasediments of latest Precambrian and Cambro-Ordovician age, the latter belonging to the so-called “Thuringian facies.” This facies is characterized by a para-autochtonous sequence of neritic and then hemipelagic rocks, documenting a marked tendency toward deeper water and interpreted as representing the passive margin of a continental unit to the northwest, which formed during an important Early Paleozoic phase of extension (e.g., Falk et al., 1995). Felsic meta-igneous rocks also occur and can be subdivided into plutonic and volcanic types, based on field and textural evidence (Siebel et al., 1997, and references therein). Three main plutonic bodies (Wunsiedel, Selb, and Waldershof) were studied for elemental and Sr-Nd isotope geochemistry by Siebel et al. (1997). These orthogneisses contain relictual granitic textures, with large, up to 10-cm-long K-feldspar megacrysts in the Wunsiedel gneiss. They were emplaced as discordant intrusions (e.g., Waldershof) in the upper crust, as shown by low-grade country rocks. This shallow emplacement level is also supported by the regional association with broadly similar, concordant formations interpreted as deriving from metavolcanic rocks on the basis of gradual contacts with surrounding sediments, and by the presence of bipyramidal or corroded quartz crystals (Siebel et al., 1997, and references therein). Their mineralogical composition is rather monotonous with quartz, K-feldspar, oligoclase, muscovite, biotite, and accessories (tourmaline, apatite, sphene, garnet, pyrite, zircon). A whole-rock Rb-Sr isochron age of 480 ± 8 Ma was measured for the Wunsiedel orthogneiss (Siebel et al., 1997). Metabasic rocks of similar age are rare in the Fichtelgebirge and correspond to alkaline basalts (Okrusch et al., 1989). Based on the peraluminous character (A/CNK > 1.08), high initial 87Sr/86Sr (>0.709), and negative εNd500 values (−3.5 to −6.5), Siebel et al. (1997) interpreted these felsic igneous rocks as dominantly crustal-derived melts and inferred an overall extensional setting and mantle heat input for melt generation. Metavolcanic Suites Kaczawa Complex (West Sudetes) The low-grade metamorphic Kaczawa complex in the west Sudetes crops out in several fault-bounded units, each composed of various parts of a Cambrian-Ordovician succession and sedimentary-tectonic mélange bodies. The lower part of the succession, exposed predominantly in the south, comprises composite metavolcanic suites and associated metasedimentary rocks (Baranowski et al., 1990). The metavolcanic rocks can be classified into three geochemical groups (Furnes et al., 1994): (1) transitional
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tholeiitic-alkaline metabasalts (pillow lavas) similar to continental rift–related magmas; (2) interlayered rhyodacitic lavas and metavolcaniclastic rocks of crustal derivation; and (3) alkaline bimodal suite of lavas and volcaniclastic rocks and alkaline metabasalts of shallow-intrusive character, also resembling initial-rift types of lavas. This part of the Kaczawa succession was interpreted to have been emplaced in a continental-rift setting during Cambrian-Ordovician times (Furnes et al., 1994; Seston et al., 2000). Preliminary U-Pb dating of zircon from a metatrachyte from that volcanic suite yielded an imprecise age of 511 ± 39 Ma (C. Pin, unpublished data). New SHRIMP zircon data (R. Kryza et al., 2005, personal commun.) confirm the ca. 500–485-Ma age of rhyodacites (associated in the field with within-plate tholeiites) and bimodal volcanics (alkali-basalts, trachytes). The upper part of the Kaczawa succession, exposed mainly in the northern and eastern parts of the area, is composed of thick, often pillowed mid-oceanic ridge–type metabasalts and associated minor Silurian graptolite black slates and cherts and, locally, conodont-bearing Devonian slates. This part is considered to represent a more evolved rift setting and deep-basin environment, possibly developed on an oceanic-type crust during SilurianDevonian times. The polygenetic Kaczawa mélanges, assigned mostly to the Upper Devonian and Lower Carboniferous, are interpreted as products of overlapping sedimentary and tectonic processes during the Variscan orogeny (Kryza et al., 2004, and references therein). The rhyodacitic metavolcanics are fine-grained and often strongly cleaved rocks composed of quartz, K-feldspar, and phengitic white mica, with minor albite, chlorite, carbonate, and scarce opaque minerals. These rocks (samples N20A, OK9.8) display rather flat trace element and REE patterns, at ~10–30 times chondrite values, with strong negative Ti and Eu anomalies (Fig. 3). The values of εNd500 for these rocks are around −3.6 (Furnes et al., 1994). The bimodal volcanic suite of the Kaczawa complex comprises alkaline lavas ranging in composition from alkali basalts to pantellerites, with trachytes predominating. They form small domes and shallow intrusions. The felsic rocks are massive, aphanitic, and mostly aphyric, with only scarce feldspar phenocrysts and matrix composed of K-feldspar, albite, quartz, phengite, titanite, and opaques. Na-amphiboles and relict jadeite were locally ascertained (Furnes et al., 1994, and references therein). In contrast to the rhyodacites, the bimodal alkaline rocks are strongly fractionated, with high contents of most incompatible trace elements. Their REE patterns are smooth, displaying strong LREE enrichment. The εNd500 values for these bimodal volcanics are different from those for the metarhyodacites and range between +1.9 and +3.7 (Furnes et al., 1994). Eastern and Southern Envelope of the Karkonosze Pluton The eastern and southern parts of the Izera-Karkonosze Massif (Fig. 1) are interpreted to comprise a nappe pile formed by northwestward overthrusting of metasedimentary-volcanic Paleozoic sequences onto the Izera-Kowary orthogneisses and their envelope
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Figure 3. Selected trace element diagrams for representative samples of ca. 500-Ma (meta)volcanic rocks from the northern part of the Bohemian Massif (normalized to chondrites, after Nakamura, 1974, and Thompson, 1982). Diagram made using GCDkit software of Janoušek et al. (2003). Data sources: Kaczawa Complex—Furnes et al. (1994); south and east Karkonosze—Kryza et al. (1995), Dostal et al. (2001); Vesser area— Bankwitz et al. (1994); Fichtelgebirge—Siebel et al. (1997); Krˇivoklát-Rokycany mafic and felsic volcanics—this study. Hbl—hornblende. The symbols of the samples are from the original papers cited.
Late Cambrian (ca. 500 Ma) felsic anorogenic magmatism in the Bohemian Massif (Mazur, 1995; Mazur and Aleksandrowski, 2001). The structurally lower nappe, the South Karkonosze unit, is composed of volcanic and sedimentary rocks affected by ca. 360-Ma HP-LT metamorphism. The upper nappe corresponds to the epidote-amphibolite– facies meta-igneous Leszczyniec complex, formed of mafic and felsic plutonic and (sub)volcanic rocks. The northwestward nappe stacking was followed by the southeast-directed Early Carboniferous extensional collapse and ca. 330-Ma intrusion (Pin et al., 1987) of the nearly undeformed Karkonosze granite. The volcanic suite of the South Karkonosze unit is represented by mafic and felsic lavas and volcaniclastics dated at 501 ± 8 Ma by the Rb-Sr whole-rock method (Bendl and Patocˇka, 1995); a broadly similar age of 485 ± 4 Ma was obtained by the U-Pb method on zircons from mafic blueschists in this area (Timmermann et al., 1999). The metavolcanics are associated with phyllites, Cambrian marbles (Hladil et al., 2003), and slates. The mafic rocks are geochemically similar to alkali basalts- to enriched tholeiites and are interpreted to have been generated in an evolved intracontinental rift setting. The felsic rocks were derived either by crustal melting or by differentiation of mantlederived magmas with substantial incorporation of crustal melts (Patocˇka et al., 1997; Dostal et al., 2001). The metaigneous Leszczyniec complex on the eastern side of the Karkonosze pluton consists of basic, intermediate, and acidic rocks, some of them dated at ca. 500 Ma by the U-Pb zircon method (Oliver et al., 1993). The mostly fine-grained mafic rocks are depleted in the most incompatible trace elements and show highly radiogenic Nd isotope signatures typical of basaltic melts extracted from strongly depleted mantle sources, such as normal-type mid-oceanic ridge basalts (N-MORBs). The felsic rocks occur as intercalations ranging from several centimeters to some hundreds of meters thick. Some of these rocks display relics of primary porphyritic textures and are interpreted to represent lavas or subvolcanic rocks. Their composition is dominated by quartz and plagioclase (<5% An), with subordinate bluishgreen amphibole, epidote, chlorite, and occasional stilpnomelane (Kryza et al., 1995). Sm-Nd data obtained on representative samples of both mafic and felsic rock-types are listed in Table 2 and discussed in combination with trace element data published by Kryza et al. (1995). Two metadiorite samples (FR28 and FR29) are enriched in LREE and Th, with clear negative Nb anomalies (Fig. 3), and have εNd500 values of +1.6 and +2.8, respectively. These plutonic rocks might simply reflect crustal contamination processes or have supra-subduction zone affinities. The latter possibility seems likely for sample FR11, which has both high Th/Nb (0.18) and high εNd500 (+7.4), a feature difficult to explain in terms of crustal contamination. Bearing in mind the very complex tectonic context of this area, which might correspond to an accretionary prism, it is not clear whether these rocks belong to the same unit as the rocks dated at ca. 500 Ma by Oliver et al. (1993). The rest of the mafic samples show affinities with NMORB, that is, depletion of Th and LREE, no significant Nb anomalies (based on X-ray fluorescence spectrometric data), and strongly radiogenic Nd isotopes (+5.6 < εNd500 < +7.9).
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Vesser Area (Northwestern Saxo-Thuringian Zone) The Vesser area (Thuringian Forest) displays the northwesternmost exposures of the Bohemian Massif. Its boundaries with basement outcrops to the south (Thuringian slate belt on the northern flank of the Schwarzburg anticline) and to the north (Ruhla Massif of the mid-German crystalline rise) are concealed by Permian deposits. The Early Paleozoic rocks consist of a ~1200-m-thick bimodal association of volcanic, volcaniclastic, and subvolcanic rocks with minor slate intercalations, strongly overprinted by deformation and greenschist-facies metamorphism (Bankwitz et al., 1994). The igneous rocks have been studied for trace elements and Nd isotopes (Bankwitz et al., 1994). More recently, the inferred Cambrian age was confirmed by U-Pb zircon dating, ranging from 513 ± 5 Ma and 508 ± 2 Ma for dacitic pyroclastics, to 503 ± 8 and 502 ± 2 Ma for microgranitic and gabbroic, respectively, high-level intrusions (Kemnitz et al., 2002). These radiometric data reverse the succession originally proposed and document a change from early rhyolitic-dacitic volcanism to a bimodal one, with locally important mafic intrusions (Kemnitz et al., 2002). The sedimentary succession shows similarities with that of uppermost Cambrian and Tremadoc of the Saxo-Thuringian realm; it began during early, mostly rhyolitic, volcanism and evolved from shallow shelf through deep shelf to pelagic environments, reflecting the progressive flooding of an older (Cadomian) subsiding basement (Kemnitz et al., 2002). Concomitantly, igneous rocks evolved toward more mafic compositions. Metabasalts and gabbroic intrusions display variable trace element features, from early, enriched (Zr/Nb ~8) to late-stage, depleted (Zr/Nb > 20) magmas. εNd values range from ~+2 or +3 in within-plate enriched basalts to +6 in gabbros, and even +7.6 in an uncontaminated tholeiitic basalt (Bankwitz et al., 1994). This trend documents a clear temporal evolution of the mantle sources involved during the igneous event. Likewise the felsic lavas show a dramatic shift of εNd values from very unradiogenic (approximately −12) in the early-stage rhyolitic volcanics, through values close to zero, to radiogenic values (+3.7) in rhyo-dacitic rocks from the bimodal association (Bankwitz et al., 1994). The combined sedimentary and igneous evidence of the
TABLE 2. Sm-Nd ISOTOPE DATA FOR THE LESZCZYNIEC META-IGNEOUS COMPLEX 147 144 143 144 Sample Rock type Sm Nd Sm/ Nd Nd/ Nd 2 εNd500Ma S.E. FR25 Orthogneiss 4.05 12.3 0.1991 0.512966 8 +6.2 FR26 Orthogneiss 4.41 17.1 0.1560 0.512844 8 +6.6 FR27 Orthogneiss 6.40 19.9 0.1943 0.512982 7 +6.8 FR11 Metavolcanic 3.11 9.24 0.2034 0.513040 8 +7.4 FR30 Metavolcanic 2.32 6.82 0.2055 0.513028 7 +7.0 FR31 Metavolcanic 1.19 3.39 0.2124 0.513010 9 +6.2 FR32 Metavolcanic 2.85 7.66 0.2248 0.513019 7 +5.6 FR33 Metavolcanic 3.21 9.77 0.1986 0.513033 8 +7.5 FR34 Metavolcanic 3.55 10.1 0.2119 0.513093 16 +7.9 Jan Metavolcanic 3.15 9.53 0.1998 0.513034 7 +7.5 FR28 Metadiorite 1.45 5.59 0.1567 0.512590 11 +1.6 FR29 Metadiorite 1.13 3.94 0.1728 0.512707 8 +2.8 Note: S.E.—standard error.
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Vesser zone provides a clear record of a Late Cambrian continental break-up event and the development of an intraplate rift, which probably evolved into the subsiding passive margin of a mature (Rheic?) ocean. Krˇivoklát-Rokycany Volcanic Complex (Barrandian Domain) By virtue of a very limited Variscan overprinting, the Barrandian realm in central Bohemia offers an excellent opportunity to study Late Proterozoic and Early- to Mid-Paleozoic sequences (Chlupácˇ et al., 1998, and references therein). In this domain, a well-defined angular unconformity separates the folded Proterozoic basement from the transgressive Cambrian or Ordovician cover. In the Prˇ ibram-Jince basin, sedimentation started in Early Cambrian times with the deposition of ~2000 m of molasse-type conglomerates and sandstones in a continental trough. This event was followed by marine transgression in the Middle Cambrian and deposition of shales (Jince Formation), then regression and partial erosion in the Late Cambrian. Subsequent deposits of Tremadocian of the Prague basin reflect the inception of a new sedimentary cycle that lasted without major break until the MidDevonian (Chlupácˇ et al., 1998). Although volcaniclastic admixtures have been recognized in many types of Lower Cambrian deposits, neither volcanics nor volcaniclastics were recorded in the Middle Cambrian deposits (Chlupácˇ et al., 1998). Major magmatism occurred in the Late Cambrian, with the effusion of two large, mostly subaerial volcanic complexes, the Strašice complex in the south, mostly composed of basic and intermediate volcanics, and the Krˇ ivoklát-Rokycany complex in the north. Only the Krˇivoklát-Rokycany complex (Waldhausrová, 1971), covering an area of ~180 km2, is dealt with in this work. It consists of a ~5-km-wide northeast–southwest-trending belt, originated from fissure-type volcanoes that produced a ≤1500-m-thick sequence of flows and ash flows of mostly acid and subordinate intermediate and mafic rocks (Chlupácˇ et al., 1998). Based on field and petrochemical evidence, Waldhausrová (1971) distinguished four eruptive stages: the volcanism started with aphanitic dacites of maximum thickness ~200 m (first group); large volumes of more mafic rock-types (basaltic andesites, andesites, and agglomerates) ~500–600 m thick were then emplaced (second group); the third group is made of smaller amounts of porphyritic dacites and rhyodacites as dome-shaped bodies; these were followed by the last group, up to 800 m thick, composed only of rhyolitic lavas (mostly porphyritic, although fluidal varieties and ignimbrites also occur). A ca. 490-Ma Rb-Sr whole-rock isochron age was obtained for the Krˇivoklát-Rokycany complex lavas, with an initial 87Sr/86Sr of 0.7041 ± 3 (Vidal et al., 1975). This age was recently revised by zircon dating (SHRIMP) to 499 ± 4 Ma (Drost et al., 2004). These authors also provided major and trace element data documenting the affinity of rhyolites with A-type granites. On the basis of REE data, they interpreted the mafic rocks as within-plate basalts similar to enriched MORB (E-MORB). These geochemical data favor generation of the Krˇivoklát-Rokycany complex in an incipient rift, probably formed in a transtensional tectonic setting (Drost et al., 2004).
ˇ IVOKLÁT-ROKYCANY TABLE 3. Sm-Nd ISOTOPE DATA FOR THE KR VOLCANIC COMPLEX 147 144 143 144 Sample Rock type Sm Nd Sm/ Nd Nd/ Nd 2 εNd500Ma TDM (Ma) S.E. KRB-1 Felsic 4.83 13.4 0.2174 0.512720 7 +0.2 – KRB-2 Intermediate 7.40 29.8 0.1503 0.512739 8 +4.9 750 KRB-3 Intermediate 7.52 30.2 0.1507 0.512746 8 +5.0 740 KRB-4a Intermediate 6.49 26.9 0.1460 0.512747 6 +5.3 690 KRB-5a Felsic 9.02 36.7 0.1376 0.512540 8 +1.8 1020 KRB-5b Felsic 5.65 18.6 0.1843 0.512622 7 +0.4 2200 KRB-6 Felsic 14.4 56.6 0.1532 0.512530 6 +0.6 1300 KRB-7 Intermediate 8.73 33.9 0.1558 0.512737 6 +4.5 830 KRB-8 Intermediate 11.0 45.7 0.1453 0.512605 6 +2.6 990 KRB-9 Intermediate 7.13 27.2 0.1586 0.512763 6 +4.8 800 KRB-10 Intermediate 6.18 24.5 0.1524 0.512743 10 +4.8 770 Notes: Analytical details as in the Appendix. S.E.—standard error; – indicates data not available.
Additional major and trace element data, and new Sm-Nd results for seven intermediate rocks and four rhyolites are given in Tables 3 and 4. The samples of intermediate bulk composition (58–65 wt% SiO2) display homogeneous patterns on chondritenormalized incompatible element diagrams (Fig. 3) with high abundances of Th (90–200 × chondrite) and large negative anomalies of Nb, Sr, and Ti. REE patterns (not shown) are gently sloping from La to Lu and parallel to those for more mafic samples (48–53 wt% SiO2) analyzed by Drost et al. (2004). A clear negative Eu anomaly, more pronounced in the most REE-rich samples, is observed in the intermediate samples only. The comparison of incompatible element patterns for mafic and intermediate samples (Fig. 3) highlights a close similarity (i.e., parallel patterns with deepening anomalies of Sr, Eu, and Ti with increasing silica content), suggesting a genetic relationship through fractionation of a mineral assemblage poor in Th and REE, but rich in Ti (i.e., mafic mineral[s], plagioclase, and Fe-Ti oxide). Interestingly, all the mafic samples have distinct negative Nb anomalies, at variance with the E-MORB affinity reported by Drost et al. (2004) on the basis of REE features alone. Rather, the combined REE and high-field-strength elements (HFSE) fingerprint is reminiscent of continental tholeiites (e.g., Dupuy and Dostal, 1984). εNd500 values obtained for intermediate rocks are all clearly positive (Table 3), implying ultimate derivation from a mantle reservoir with time-integrated depletion of LREE. With one exception (KRB-8, with εNd500 = +2.6), Nd isotopes show little variation over a range of silica content (58.5–65.3 wt%), and the most radiogenic value (+5.3) is measured in the most silicic sample (KRB-4a, with 65.3 wt% SiO2). This observation is not in line with a simple assimilation–fractional crystallization (AFC) model to explain the isotopic variation within intermediate rocks, but rather suggests that heterogeneous sources and/or different evolutionary paths of magma batches were involved. The rhyolites have much lower, albeit still positive, εNd500 values, from +0.2 to +1.8. These values preclude derivation from ancient crustal rocks with time-integrated LREE enrichment, in keeping with the relatively low Sr initial ratio (~0.704; Vidal et al., 1975). Our data are still too limited to draw firm inferences on the petrogenesis of these rhyolites. Samples KRB-8 (65 wt% SiO2, εNd500
Late Cambrian (ca. 500 Ma) felsic anorogenic magmatism in the Bohemian Massif
Sample (oxide wt%) SiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O TiO2 P2O5 LOI Total (ppm) Nb Sr Th Y Zr
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TABLE 4. CHEMICAL ANALYSES OF MAJOR ELEMENTS AND TRACE ELEMENTS OF VOLCANIC ROCKS FROM THE KRˇVOKLÁT-ROKYCANY COMPLEX (BARRANDIAN AREA) KRB-7 KRB-10 KRB-9 KRB-2 KRB-3 KRB-8 KRB-4a KRB-5a KRB-5b KRB-6 57.64 16.08 8.88 0.145 1.23 4.90 4.04 1.65 1.04 0.401 4.23 100.23
58.53 16.80 7.89 0.107 1.89 4.07 4.70 0.99 0.811 0.246 4.58 100.60
60.32 17.09 7.50 0.067 1.91 2.92 4.38 1.94 0.827 0.313 2.85 100.10
9.36 156 5.22 54.5 342
6.32 187 4.18 41.0 276
7.06 163 4.54 45.8 293
60.76 16.77 6.76 0.054 1.05 1.77 7.03 1.20 0.634 0.336 3.05 99.41 8.16 124 4.48 43.0 323
61.41 16.93 7.31 0.046 1.08 1.59 7.09 1.09 0.651 0.349 2.98 100.53
65.15 14.92 7.54 0.127 0.837 2.67 4.21 3.01 0.694 0.265 1.56 100.97
65.27 16.17 5.71 0.091 1.17 1.24 5.95 1.53 0.551 0.232 2.28 100.18
74.04 13.11 2.51 0.026 0.274 0.135 2.92 4.98 0.201 0.04 2.46 100.69
74.15 14.15 1.89 0.018 0.263 0.134 7.37 0.42 0.102 0.036 1.56 100.09
8.09 107 4.39 43.0 317
12.0 122 9.02 64.2 802
6.65 226 4.23 38.8 287
11.6 32.2 13.4 73.6 330
7.38 25.1 12.2 60.3 155
74.48 12.58 1.70 0.01 0.103 0.096 2.60 5.35 0.126 0.038 1.34 98.42 16.1 14.3 15.7 147 230
KRB-1 75.52 13.12 0.697 0.001 0.178 0.056 0.625 7.48 0.113 0.07 1.71 99.57 8.44 25.2 5.98 81.9 93.6
La 23.0 17.9 17.2 19.3 18.9 33.8 20.0 34.1 14.4 46.5 9.04 Ce 55.3 42.1 42.2 45.7 45.2 80.8 46.7 77.8 33.3 113,000 23.0 Pr 7.60 5.66 5.89 6.64 6.61 10.7 6.18 9.74 4.43 13.8 3.29 Nd 33.3 24.6 26.6 28.8 28.7 44.8 26.2 38.5 18.5 56.2 13.3 Sm 8.75 6.33 7.08 7.16 7.15 11.0 6.41 9.31 5.80 14.4 6.12 Eu 1.93 1.45 1.70 1.74 1.60 1.97 1.60 0.729 0.412 0.961 0.317 Gd 9.40 6.70 7.63 7.38 7.25 11.1 6.57 9.98 7.39 17.3 10.5 Tb 1.50 1.08 1.25 1.22 1.20 1.80 1.06 1.86 1.46 3.24 2.30 Dy 9.35 6.77 7.79 7.43 7.35 11.0 6.56 12.2 9.87 21.7 14.6 Ho 1.89 1.41 1.59 1.49 1.48 2.24 1.32 2.50 2.05 4.57 2.62 Er 5.39 4.04 4.57 4.36 4.30 6.48 3.80 7.36 6.07 13.1 6.57 Tm 0.793 0.618 0.681 0.651 0.651 0.971 0.569 1.12 0.917 1.95 0.876 Yb 5.23 4.08 4.53 4.42 4.50 6.60 3.82 7.48 6.18 12.6 5.31 Lu 0.818 0.644 0.710 0.688 0.695 1.05 0.618 1.15 0.941 1.89 0.746 Note: The major and trace element data were obtained at the Centre de Recherche Pétrogaphique et Géochemique, Nancy, France, by inductively coupled plasma (ICP) atomic emission spectrometry and ICP–mass spectrometry, respectively, following methods described by Carignan et al. (2001). LOI—loss on ignition.
= +2.6) and KRB-5a (74 wt% SiO2, εNd500 = +1.8) suggest that at least some of the rhyolites might have evolved from mantlederived, mafic to intermediate parental magmas, through AFC in a crustal magma chamber. Alternatively, the rhyolites or their precursor might have been produced in the lower crust, through magma mixing between mantle and crustal end-members, or generated directly through partial melting of sources that were characterized by Nd isotopes intermediate between those of depleted mantle and typical continental crust. Based on the geochemical affinity of rhyolites with A-type granitoids (Drost et al., 2004), possible crustal sources might include igneous rocks of tonalitic to granodiorite composition (e.g., Creaser et al., 1991) or partially dehydroxylated biotite- and amphibole-bearing gneisses (e.g., Skjerlie and Johnston, 1992). Provided they occurred at that time in the Barrandian lower crust, Neoproterozoic graywackes (e.g., Jakeš et al., 1979) might provide suitable source materials, as suggested by preliminary Nd isotope data (εNd500 ≈ −1). In this scenario, under- and intraplating by basaltic magmas similar to continental tholeiites (as represented by early-stage and
volumetrically subordinate mafic rocks in the Krˇ ivoklát-Rokycany complex) would have caused HT metamorphism and partial dehydroxylation of the Barrandian lower crust, followed by dehydration-melting to produce rhyolitic magmas. The AFC and partial melting models are not mutually exclusive and might have operated during the same igneous event. DISCUSSION Main Features and Inferred Source Materials Orthogneisses Characteristically, the numerous occurrences of ca. 500-Ma gneisses from the northern part of the Bohemian Massif consist of leucocratic rocks that display two main varieties: (1) coarseto medium-grained rocks with K-feldspar megacrysts (augen gneisses; e.g., S´niez˙nik gneisses, most Izera metagranitoids, some Red Gneisses from eastern Erzgebirge, Wunsiedel gneisses); and (2) finer-grained, equigranular rocks, often more intensely
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deformed (e.g., Gierałtów gneisses, some Izera metagranites, some Red Gneisses from Erzgebirge), but transitional textural varieties are also observed, at least in part caused by heterogeneous ductile deformation. The emplacement level of igneous precursors is often unconstrained because of the intense tectonometamorphic overprinting, having reached eclogite- or granulite-facies conditions in several cases (Gierałtów, Góry Sowie, Red Gneisses). However, relatively shallow levels of intrusion can be inferred in certain places (e.g., Izera, Waldershof gneiss) where Variscan deformation and metamorphism were limited, in general, to low-grade conditions. Although variable in terms of textures (porphyritic versus equigranular) and degree of ductile deformation, all these metagranitoids are leucocratic, muscovitebearing rocks that share an important common feature, namely, the absence of dark, microgranular enclaves. Importantly, they are not associated with intermediate (dioritic) or mafic lithologies, except rare metabasite bodies (sometimes eclogitized) interpreted to reflect boudinaged basaltic dikes emplaced after—and unrelated to—granite magmatism. The mineralogical composition of the gneisses is broadly homogeneous and monotonous with quartz, K-feldspar, relatively sodic plagioclase, muscovite, and biotite, documenting their origin as two-mica granites. In most occurrences (except the Izera gneisses, for which prismatic cordierite is widespread), highly aluminous minerals (cordierite, garnet, sillimanite) are scarce or even absent, only primary muscovite reflecting slightly peraluminous igneous protoliths. The major element composition of these granitoids (e.g., Borkowska et al., 1990; Kröner et al., 1995; Siebel et al., 1997; Turniak et al., 2000; Mingram et al., 2004; Oberc-Dziedzic et al., 2005a,b) shows high contents of silica (68–77 wt%), alumina (12–15 wt%), and alkali elements (7–9 wt% Na2O + K2O); low concentrations of Ca, Mg, and Fe; and a variably pronounced peraluminous character (A/CNK > 1). Although some mobility of large-ion lithophile elements certainly occurred during deformation and metamorphism, implying that these elements should be used with caution, K2O > Na2O in all cases, and CaO contents rarely exceed 1.5 wt%. Overall, these rocks are reminiscent of Stype granitoids (e.g., Clemens, 2003, and references therein). Chondrite-normalized patterns for selected trace elements considered as relatively immobile (except Sr) during deformation and metamorphism (Fig. 2) display similar features in all examples studied, specifically, strong enrichment in Th (typically 100–400 × chondrite abundance) and LREE compared to middle REE and HREE. Typically, the LREE abundances decrease with increasing silica content, or with increasing Rb/Sr ratio used as a differentiation index. Concomitantly, the negative Eu anomaly deepens markedly, and the degree of fractionation of the HREE decreases, as shown by the transition from steep to almost flat chondrite-normalized HREE patterns in the more evolved samples (e.g., Fichtelgebirge orthogneisses; Fig. 2). This difference may be interpreted to reflect the fractionation, during magma differentiation, of a mineral assemblage containing plagioclase and one or several accessory phases rich in LREE and middle REE relative to HREE. In some cases (e.g., Izera, Wunsiedel),
the degree of LREE enrichment decreases in the most evolved samples, giving flat to LREE-depleted patterns, probably as a result of monazite fractionation. Deep negative anomalies of Nb, Ti, Sr, and Eu occur throughout. Although crystal fractionation (e.g., plagioclase for Sr and Eu) certainly played a role, these trace element data are interpreted to reflect derivation from source materials that were themselves enriched in Th and LREE, and relatively depleted in Sr and elements of the Ti group, as typical for granitoids, quartzofeldspathic gneisses, pelitic and graywackeous mature sediments, and crustal materials in general (e.g., Taylor and McLennan, 1985). The distinct Ti-Nb negative anomalies and HREE fractionation shown by the least evolved samples further suggest that the primary magmas separated from a rutile- and garnet-bearing restitic assemblage, as expected for granulitic residues left behind after partial melting of most metasedimentary sources. It is likely, however, that the negative Nb anomalies observed in the 500-Ma granitoids were in great part inherited from their crustal source materials. Radiogenic isotope data provide important, albeit not unequivocal, constraints on the possible source materials involved. Although late-stage disturbances of the Rb-Sr system occurred in many cases during metamorphism and deformation, whole-rock Rb-Sr isochrons broadly preserved igneous emplacement ages in some occurrences, as shown by reasonably good agreement with U-Pb zircon data. In these cases, initial 87Sr/86Sr ratios are fairly radiogenic (0.715 ± 5 for Gierałtów gneisses, Borkowska et al., 1990; 0.709 ± 1 for Izera granitoids, Borkowska et al., 1980; 0.7095 ± 7 for the Wunsiedel orthogneiss, Siebel et al., 1997) and point to source materials with relatively high time-integrated Rb/ Sr ratios, such as relatively mature upper crustal sediments. This possibility is further corroborated by largely negative εNd values (Fig. 4) throughout (mean −5.0, standard deviation [SD] 1.0 for seventeen samples from the Orlica-S´niez˙nik dome, Kröner et al., 2001, and this study; approximately −6 for two samples from Góry Sowie, Kröner and Hegner, 1998; mean −5.5, SD 0.7 for thirteen samples from Izera, Kröner et al., 2001, and ObercDziedzic et al., 2005a; mean −5.1, SD 1.8 for five samples from Erzgebirge Red Gneisses, Kröner et al., 1995; mean −3.5, SD 0.3 for seven samples from Fichtelgebirge, Siebel et al., 1997). These values reflect extraction from source reservoirs enriched in LREE (i.e., with low Sm/Nd) on a time-integrated basis. Average crustal residence ages have been calculated (or recalculated, for the sake of homogeneity) relative to the island-arc type depleted mantle model of DePaolo (1981). Samples with 147Sm/144Nd > 0.15 were discarded from these calculations, to keep to a minimum the effect of late-stage increases in the Sm/Nd ratio, probably caused by monazite fractionation, which would lead to spuriously old model ages. The results range from 1.2 to 2.1 Ga, with a majority of values falling between 1.5 and 1.8 Ga. Bearing in mind the widely acknowledged S-type character of most 500-Ma granitoids (e.g., Siebel et al., 1997; Turniak et al., 2000; Tichomirowa et al., 2001; Mingram et al., 2004; Oberc-Dziedzic et al., 2005a,b), these dates should not be taken as evidence, even circumstantial, for derivation from Mid-Proterozoic crust, and more especially,
Late Cambrian (ca. 500 Ma) felsic anorogenic magmatism in the Bohemian Massif
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x Góry Sowie orthogneisses Góry Sowie migmatites ´ niez˙ nik Massif - S´niez˙ nik gneiss Orlica-S ´ niez˙ nik Massif - Gieraltów gneiss Orlica-S Izera orthogneisses
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Figure 4. εNd vs. Sm/Nd data compiled for ca. 500-Ma (A) orthogneisses and (B) (meta)volcanic rocks from the northern part of the Bohemian Massif. Data sources: see Figures 2 and 3, and this study. Ca. 550-Ma gneisses from Erzgebirge analyzed by Kröner et al. (1995) and Tichomirowa et al. (2001) are included for comparison (with εNd recalculated for 500 Ma). E-KAR—east Karkonosze; ERZ—Erzgebirge; FICH— Fichtelgebirge; GIE—Gierałtów; GSM—Góry Sowie Massif; Hbl—hornblende; IZE—Izera Massif; KC-RD—Kaczawa-rhyodacites; KC-T— Kaczawa-trachytes; KR-I—Krˇivoklát-Rokycany-intermediate rocks; KR-R—Krˇivoklát-Rokycany-rhyolites; Ms—muscovite; S-KAR—south Karkonosze; SNI—S´niez˙nik; VES—Vesser area.
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as a reason to infer the existence of any ancient basement at depth. Instead of referring to specific crustal formation event(s), the TDM ages most likely reflect the average of several detrital components mixed in the granitoid sources, insofar as no significant mantle material was added at the time of magma generation. This interpretation is supported by complex age spectra of inherited zircons that occur frequently in these rocks. For example, a 2.1-Ga inherited component was documented in the Rumburk granite (Oliver et al., 1993), and zircon grains up to 2.5–2.6 Ga have been found elsewhere (e.g., Kröner et al., 1995; Turniak et al., 2000; Kryza and Fanning, 2007). These point to the presence of old, most likely multiply recycled, crustal components in the source of the ca. 500-Ma granitoids. Geologically more interesting is the occurrence, documented by SHRIMP measurements (Turniak et al., 2000), and possibly, by Pb/Pb evaporation ages (Kröner et al., 1995), of inherited igneous zircons as young as ca. 540 Ma in some cases. These zircons document the young depositional age (and/or intrusion age) of the volcaniclastic sediments (and/or S-type granitoids), which formed the probable source materials of the 500-Ma orthogneisses. In summary, the combined radiogenic isotope data suggest that most of these granitoids correspond to partial melts of heterogeneous sedimentary sources containing both ancient (2 Ga and/or older) mature, recycled components, and recent erosion products from Late Proterozoic igneous sources. The relative contribution of these two contrasting sedimentary components to the granitoid source material is likely to vary from place to place. For example, the old recycled source (metapelites?) probably played an important role in the genesis of the strongly peraluminous Izera granitoids, with εNd values of approximately −5 to –7, and TDM in the range of 1.73–2.17 Ga (Oberc-Dziedzic et al., 2005a). In contrast, less mature sediments, such as graywackes containing a substantial igneous component of Late Proterozoic age in addition to old recycled epiclastic materials, were probably more important in the source of the Wunsiedel orthogneiss (εNd approximately −3; Siebel et al., 1997) or some of the Red Gneisses (εNd approximately −3 to –4; Kröner et al., 1995). Bearing in mind the possible disturbances of Si, Na, K, and Ca during Variscan deformation and metamorphism and the likelihood of some restite entrainment (Al, Fe, Mg) in anatectic melts, the major element composition of the ca. 500-Ma orthogneisses fits rather well with the range of melts experimentally produced from most common crustal rocks in the absence of fluids. Indeed, except under special circumstances, the amount of free water available in the lower crust is expected to be too low to make H2O-saturated melting a significant process (e.g., Thompson and Connolly, 1995; Clemens and Watkins, 2001). In broad terms, detrital sedimentary rocks, such as metapelites and metapsammites (graywackes and quartzofeldspathic gneisses), are the most fertile protoliths for dehydration melting, because they contain appropriate relative proportions of quartz and feldspar on the one hand, providing a low-temperature melting fraction, and micas on the other hand, providing H2O for lowering the melting temperature (Thompson, 1996, and references therein). Although
potentially fertile in terms of quartzofeldspathic components, common granites and granodiorites are too anhydrous (i.e., do not contain enough biotite + muscovite) to provide much granitoid magma through remelting (e.g., Patiño Douce and Johnston, 1991). Strongly peraluminous granitic melts could be produced from amphibolite protoliths (Patiño Douce and Beard, 1995), but such source rocks are considered rather unlikely in the present case, based on high potassium contents, radiogenic Sr isotopes, largely negative εNd values, and the widespread occurrence of inherited zircons in the studied orthogneisses, which instead favor a metasedimentary source. According to Thompson (1996), the interpretation of experimental results indicates that pelites (containing more mica than feldspar and quartz) would be more fertile than psammites (containing more feldspar and quartz than micas) at low P (e.g., 5 kilobars) and when the plagioclase is calcic (e.g., An40–An50). In contrast, psammites would be more melt productive than pelites at higher pressure and when plagioclase is more sodic. For these reasons, it is considered that the optimum protolith for lower crustal anatexis is close to a metagraywacke (e.g., Clemens and Vielzeuf, 1987; Patiño Douce and Johnston, 1991; Thompson, 1996), because of its large plagioclase (providing Na2O) and quartz contents, and relative scarcity of aluminosilicate. Together with geochemical and isotope data, these considerations give credence to a scenario explaining the generation, ca. 500 Ma ago, of large volumes of granitic magmas through partial melting of highly fertile metasedimentary materials, such as the Neoproterozoic graywackes, which form a significant proportion of the outcrops of the northern Bohemian Massif. Volcanic Suites In contrast to the broad compositional and isotopic homogeneity of the orthogneisses, the coeval felsic volcanics or hypabyssal intrusions show a great range of elemental and isotopic variation. At one extreme, some rhyolites or dacites display extremely nonradiogenic Nd isotopes, which provide evidence for interpreting them as almost pure crustal melt. The typical example for that group is provided by an early-stage, rhyolitic lava from the Vesser area (sample 7904 from Bankwitz et al., 1994) with an initial εNd value as low as −11.7. A marked depletion of the HREE (<10× chondrite abundance) suggests that the parental magma separated from a garnet-rich residuum in the lower crust. The TDM model age calculated for this sample (1.9 Ga) suggests that its parental magma might have been generated through partial melting of ca. 2-Ga basement rocks, such as those found as small inliers in northern Armorican Massif (Icartian orthogneisses), or from sedimentary detritus eroded from a similar source, without significant addition of younger components. More widespread are felsic lavas with less extreme, but still strongly negative, εNd500 values—for example, volcaniclastic samples with εNd500 of approximately –5 in the S´wierzawa unit, and with εNd500 of approximately −3.7 in the Bolków unit, both from the Kaczawa Mountains (Furnes et al., 1994). Felsic metavolcanics (porphyroids) from the east Karkonosze complex—with εNd500 values from –4.3 to –5.5 and TDM ranging
Late Cambrian (ca. 500 Ma) felsic anorogenic magmatism in the Bohemian Massif from 1.5 to 1.8 Ga (Dostal et al., 2001)—also belong to this predominantly crustally derived group, although the relatively low 87Sr/86Sr initial ratios measured for these rocks (Bendl and Patocˇka, 1995) suggest that the Sr budget in the parental magma was not dominated by an ancient, mature component with a high time-integrated Rb/Sr ratio. The concordant gneisses from the Fichtelgebirge, interpreted as metavolcanics, can be ascribed to the same group, based on their distinctly negative εNd500 values (from –3.8 to –6.4, mean –4.9, SD 1.0, for seven samples), and model ages ranging from 1.5 to 1.8 Ga (Siebel et al., 1997). At the other extreme are rhyolites and trachytes, closely associated in the field with volumetrically predominant metabasalts of various types. In the southern Kaczawa Mountains and Vesser area, felsic lavas, such as sample 7525 from Vesser (Bankwitz et al., 1994) and samples Swie 4, P4.1, C14.15, and LA2 from Kaczawa (Furnes et al., 1994), are associated with metabasalts similar to ocean island basalts and alkali basalts. Based on elevated concentrations of LREE, Zr (600–1200 ppm), and Nb (>100 ppm), these lavas are similar to alkaline rhyolites and trachytes. They share the same Nd isotope signature (specifically, εNd500 of approximately +2 to +4) as the associated enriched basalts and are interpreted to represent evolved liquids left after closed-system differentiation of the concomitant within-plate basalts. The fractionation of HREE in these alkaline rhyolites mirrors that of the associated basalts. Conspicuous negative anomalies of Sr, Eu, and Ti suggest that low-pressure fractional crystallization of plagioclase + Fe-Ti oxide + mafic phase(s) was probably involved in this differentiation process, leading to a general increase of REE concentrations. In the Leszczyniec unit of the east Karkonosze complex, felsic rocks of plagiogranitic mineralogical composition occur in close association with Th- and LREE-depleted mafic rocks strongly reminiscent of N-MORB (Kryza et al., 1995). Geochemically, the felsic rocks are characterized by low potassium contents and almost flat chondrite-normalized patterns of incompatible elements, with the exception of Sr, Eu, Nb, and Ti, which show deep anomalies. Sm-Nd analyses (Table 2) of these samples yield very radiogenic isotope signatures (εNd500 from +6.2 to +6.8) within the range of values measured for local mafic to intermediate volcanics (from +5.6 to +7.9). Based on these data, the Leszczyniec felsic rocks resemble oceanic plagiogranites, generated either by fractional crystallization of a basaltic parent, or by low-degree melting of a similar protolith. Deciphering the specific petrogenesis of these rocks is beyond the scope of this article, but the occurrence of a few intermediate samples (FR11 and FR31) with 55 and 64 wt% SiO2 and εNd500 +7.4 and +6.2, respectively, might lend some support to the fractional crystallization model. Besides those felsic rocks which clearly prove to be either largely of crustal origin (as shown by strongly negative εNd) or genetically linked with the enriched, within-plate type (εNd approximately +3), or depleted (εNd > +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
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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,
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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).
REFERENCES CITED ˙ elaz´niewicz, A., 1998, Północna cze˛s´c´ bloku karkoAchramowicz, S., and Z nosko-izerskiego: zapis dwóch orogenez: Polskie Towarzystwo Mineralogiczne, Prace Specjalne, v. 11, p. 9–13. Anderson, D.L., 2000, The thermal state of the upper mantle; no role for mantle plumes: Geophysical Research Letters, v. 27, p. 3623–3626, doi: 10.1029/2000GL011533. Anderson, J.L., and Bender, E.E., 1989, Nature and origin of Proterozoic Atype granitic magmatism in the southwestern United States of America: Lithos, v. 23, p. 19–52, doi: 10.1016/0024-4937(89)90021-2. Anderson, J.L., and Thomas, W.M., 1985, Proterozoic anorogenic two-mica granites: Silver Plume and St. Vrain batholiths of Colorado: Geology, v. 13, p. 177–180, doi: 10.1130/0091-7613(1985)13<177:PATGSP>2.0.CO;2. Atherton, M.P., and Ghani, A.A., 2002, Slab break-off: A model for Caledonian, Late Granite syn-collisional magmatism in the orthotectonic (metamorphic) zone of Scotland and Donegal, Ireland: Lithos, v. 62, p. 65–85, doi: 10.1016/S0024-4937(02)00111-1. Bankwitz, P., Bankwitz, E., Kramer, W., and Pin, C., 1994, Early Paleozoic bimodal volcanism in the Vesser area, Thuringian forest, eastern Germany: Zentralblatt für Geologie und Paläontologie, 1992, vol. I, no. 9–10, p. 1113–1132. Baranowski, Z., Haydukiewicz, A., Kryza, R., Lorenc, S., Solecki, A., and Urbanek, Z., 1990, Outline of the geology of the Góry Kaczawskie (Sudetes, Poland): Neues Jahrbuch für Geologie und Paläontologie. Abhandlungen, v. 179, p. 223–257. Bendl, J., and Patocˇka, F., 1995, The Rb-Sr isotope geochemistry of the metamorphosed bimodal volcanic association of the Rýchory Mts. crystalline complex, west Sudetes, Bohemian Massif: Geologia Sudetica, v. 29, p. 3–18. Berg, G., 1923, Die Gesteine des Isergebirges: Jahrbuch der Preussischen Geologischen Landes-Anstalt, Berlin, v. 43, p. 125–168. Borkowska, M., Hameurt, J., and Vidal, P., 1980, Origin and age of Izera gneisses and Rumburk granites in the western Sudetes: Acta Geologica Polonica, v. 30, p. 121–146. Borkowska, M., Choukroume, P., Hameurt, J., and Martineau, F., 1990, A geochemical investigation of the age, significance and structural evolution of the Caledonian-Variscan granite gneisses of the S´niez˙nik metamorphic area: Geologia Sudetica, v. 25, p. 1–27. Bröcker, M., and Klemd, R., 1996, Ultrahigh-pressure metamorphism in the S´niez˙nik Mountains (Sudetes, Poland): P-T constraints and geological implications: Journal of Geology, v. 104, p. 417–433. Carignan, J., Hild, P., Mevelle, G., Morel, J., and Yeghicheyan, D., 2001, Routine analyses of trace elements in geological samples using flow injection and low-pressure on-line liquid chromatography coupled to ICP-MS: A study of geochemical reference materials BR, DR-N, UB-N, AN-G and GH: Geostandards Newsletters, v. 25, p. 187–198. 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. Clemens, J.D., 2003, S-type granitic magmas—Petrogenetic issues, models and evidence: Earth Sciences Review, v. 61, p. 1–18, doi: 10.1016/S00128252(02)00107-1. Clemens, J.D., and Vielzeuf, D., 1987, Constraints on melting and magma production in the crust: Earth and Planetary Science Letters, v. 86, p. 287– 306, doi: 10.1016/0012-821X(87)90227-5. Clemens, J.D., and Watkins, J.M., 2001, The fluid regime of high-temperature metamorphism during granitoid magma genesis: Contributions to Mineralogy and Petrology, v. 140, p. 600–606.
227
Collins, W.J., Beams, S.D., White, A.J.R., and Chappell, B.W., 1982, Nature and origin of A-type granites with particular reference to south-eastern Australia: Contributions to Mineralogy and Petrology, v. 80, p. 189–200, doi: 10.1007/BF00374895. Creaser, R.A., Price, R.C., and Wormald, R.J., 1991, A-type granites revisited: Assessment of a residual-source model: Geology, v. 19, p. 163–166, doi: 10.1130/0091-7613(1991)019<0163:ATGRAO>2.3.CO;2. Crowley, Q.G., Floyd, P.A., Winchester, J.A., Franke, W., and Holland, J.G., 2000, Early Palaeozoic rift-related magmatism in Variscan Europe: Fragmentation of the Armorican Terrane Assemblage: Terra Nova, v. 12, p. 171–180, doi: 10.1046/j.1365-3121.2000.00290.x. Crowley, Q.G., Timmermann, H., Noble, S.R., and Holland, J.G., 2002, Palaeozoic terrane amalgamation in central Europe: A REE and Sm-Nd isotope study of the pre-Variscan basement, NE Bohemian Massif, in Winchester, J.A., Pharaoh, T.C., and Verniers J. eds., Palaeozoic amalgamation of central Europe. London, Geological Society of London Special Publication 201, p. 157–176. DePaolo, D.J., 1981, Neodymium isotopes in the Colorado Front Range and crust—Mantle evolution in the Proterozoic: Nature, v. 291, p. 193–196, doi: 10.1038/291193a0. Don, J., Dumicz, M., Wojciechowska, I., and Z˙ elaz´niewicz, A., 1990, Lithology and tectonics of the Orlica-S´niez˙nik Dome, Sudetes: Recent state of knowledge: Neues Jahrbuch für Geologie und Paläontologie Abhandlungen, v. 179, p. 159–188. Dörr, W., Fiala, J., Vejnar, Z., and Zulauf, G., 1998, U-Pb zircon ages and structural development of metagranitoids of the Teplá crystalline complex: Evidence for pervasive Cambrian plutonism within the Bohemian Massif (Czech Republic): Geologische Rundschau, v. 87, p. 135–149, doi: 10.1007/s005310050195. Dörr, W., Zulauf, G., Fiala, J., Franke, W., and Vejnar, Z., 2002, Neoproterozoic to Early Cambrian history of an active plate margin in the Teplá-Barrandian unit—A correlation of U-Pb isotopic dilution TIMS ages (Bohemia, Czech Republic): Tectonophysics, v. 352, p. 65–85, doi: 10.1016/S00401951(02)00189-0. Dostal, J., Patocˇka, F., and Pin, C., 2001, Middle/Late Cambrian intracontinental rifting in the central west Sudetes, NE Bohemian Massif (Czech Republic): Geochemistry and petrogenesis of the bimodal metavolcanic rocks: Geological Journal, v. 36, p. 1–17, doi: 10.1002/gj.872. Drost, K., Linnemann, U., McNaughton, N., Fatka, O., Kraft, P., Gehmlich, M., Tonk, C., and Marek, J., 2004, New data on the Neoproterozoic-Cambrian geotectonic setting of the Teplá-Barrandian volcano-sedimentary successions: Geochemistry, U-Pb zircon ages, and provenance (Bohemian Massif, Czech Republic): International Journal of Earth Sciences, v. 93, p. 742–757, doi: 10.1007/s00531-004-0416-5. Dupuy, C., and Dostal, J., 1984, Trace element geochemistry of some continental tholeiites: Earth and Planetary Science Letters, v. 67, p. 61–69, doi: 10.1016/0012-821X(84)90038-4. Eby, G.N., 1992, Chemical subdivision of the A-type granitoids: Petrogenetic and tectonic implications: Geology, v. 20, p. 641–644, doi: 10.1130/00917613(1992)020<0641:CSOTAT>2.3.CO;2. Falk, F., Franke, W., and Kurze, M., 1995, Saxothuringian basin: Stratigraphy, in Dallmeyer, R.D., Franke, W., and Weber, K., eds., Pre-Permian geology of central and eastern Europe: Berlin, Springer, p. 221–234. Favela, J., and Anderson, D.L., 2000, Extensional tectonism and global volcanism, in Boschi, E., Ekstrom, G., and Morelli, A., eds., Problems in geophysics for the new millennium: Bologna, Editrice Compositori, p. 463–498. Finger, F., Broska, I., Haunschmidt, B., Hrasko, L., Kohut, M., Krenn, E., Petrik, I., Riegler, G., and Uher, P., 2003, Electron-microprobe dating of monazites from western Carpathian basement granitoids: Plutonic evidence for an important Permian rifting event subsequent to Variscan crustal anatexis: International Journal of Earth Sciences (Geologische Rundschau), v. 92, p. 86–98. Floyd, P.A., Winchester, J.A., Seston, R., Kryza, R., and Crowley, Q.G., 2000, Review of geochemical variation in Lower Palaeozoic metabasites from the NE Bohemian Massif: Intracratonic rifting and plume-ridge interaction, in Franke, W., Haak, V., Oncken, O., and Tanner, D., eds., Orogenic processes: Quantification and modelling in the Variscan Belt: London, Geological Society of London Special Publication 179, p. 155–174. Furnes, H., Kryza, R., Muszyn´ski, A., Pin, C., and Garmann, I.B., 1994, Geochemical evidence for progressive, rift-related of early Palaeozoic volcanism in the eastern Sudetes: Journal of the Geological Society of London, v. 151, p. 91–110.
228
Pin et al.
Hanson, R.E., and Al-Shaieb, Z., 1980, Voluminous subalkaline silicic magmas related to intracontinental rifting in the southern Oklahoma aulacogen: Geology, v. 8, p. 180–184, doi: 10.1130/0091-7613(1980)8<180: VSSMRT>2.0.CO;2. Hladil, J., Patocˇka, F., Kachlik, V., Melichar, R., and Hubacˇik, M., 2003, Metamorphosed carbonates of Krkonoše Mountains and Paleozoic evolution of Sudetic terranes (NE Bohemia, Czech Republic): Geologia Carpathica, v. 54, p. 281–297. Jakeš, P., Zoubek, J., Zoubkova, J., and Franke, W., 1979, Graywackes and metagraywackes of the Teplá-Barrandian Proterozoic area: Sborník Geologických Véd, v. 33, p. 83–122. Janoušek, V., Farrow, C.M., and Erban, V., 2006, GCDkit: Interpretation of whole-rock geochemical data in igneous geochemistry: Introducing geochemical data toolkit (GCDkit): Journal of Petrology, v. 47, no. 6, p. 1255–1259. Kemnitz, H., Romer, R.L., and Oncken, O., 2002, Gondwana break-up and the northern margin of the Saxothuringian belt (Variscides of central Europe): International Journal of Earth Sciences, v. 91, p. 246–259, doi: 10.1007/ s005310100209. Korytowski, A., Dörr, W., and Z˙ elaz´niewicz, A., 1993, U-Pb dating of (meta)granitoids in the NW Sudetes (Poland) and their bearing on tectono-stratigraphic correlation: EUG 7 Meeting, Strasbourg, 4–8 April: Terra Nova v. 5, abstract supplement 1, p. 331. Kröner, A., and Hegner, E., 1998, Geochemistry, single zircon ages, and Sm-Nd systematics of granitoid rocks from the Góry Sowie (Owl) Mountains, Polish west Sudetes: Evidence for early Palaeozoic arc-related plutonism: Journal of the Geological Society of London, v. 155, p. 711–724. Kröner, A., and Mazur, S., 2003, Proterozoic and Palaeozoic crustal components across the east/central Sudetes boundary at the eastern margin of the Bohemian Massif: New U/Pb and Pb/Pb single zircon ages from the eastern fore-Sudetic block (SW Poland): Journal of the Czech Geological Society, v. 48, no. 1–2, p. 83–84. Kröner, A., and Willner, A.P., 1998, Time of formation and peak of Variscan HP-HT metamorphism of quartz-feldspar rocks in the central Erzgebirge, Saxony, Germany: Contributions to Mineralogy and Petrology, v. 132, p. 1–20. Kröner, A., Hegner, E., Hammer, J., Haase, G., Bielicki, K.-H., Krauss, M., and Eidam, J., 1994, Geochronology and Nd-Sr systematics of Lusatian granitoids: Significance for the evolution of the Variscan orogen in east-central Europe: Geologische Rundschau, v. 83, p. 357–376. Kröner, A., Willner, A.P., Hegner, E., Frischbutter, A., Hofmann, J., and Bergner, R., 1995, Latest Precambrian (Cadomian) zircon ages, Nd isotopic systematics and P-T evolution of granitoid orthogneisses of the Erzgebirge: Saxony and Czech Republic: Geologische Rundschau, v. 84, p. 437–456, doi: 10.1007/s005310050016. Kröner, A., Jaeckel, P., Hegner, E., and Opletal, M., 2001, Single zircon ages and whole rock Nd isotopic systematics of early Palaeozoic granitoid gneisses from the Czech and Polish Sudetes (Jizerské hory, Krkonoše Mts. and Orlice-Snžnik complex): International Journal of Earth Sciences, v. 90, p. 304–324, doi: 10.1007/s005310000139. Kryza, R., 1981, Migmatization in gneiss of northern part of the Sowie Góry: Geologia Sudetica, v. 16, p. 7–100. Kryza, R., and Fanning, C.M., 2007, Devonian deep-crustal metamorphism and exhumation in the Variscan orogen: Evidence from SHRIMP zircon ages from the HT-HP granulites and migmatites of the Góry Sowie (Polish Sudetes): Geodinamica Acta (in press). Kryza, R., and Pin, C., 1997, Cambrian/Ordovician magmatism in the Polish Sudetes: No evidence for subduction-related setting: EUG 9 Meeting, Strasbourg, 23–27 March: Terra Nova, v. 9, abstract supplement 1, p. 144. Kryza, R., and Pin, C., 2002, Mafic rocks in a deep crustal segment of the Variscides (the Góry Sowie, SW Poland): Evidence for crustal contamination in an extensional setting: International Journal of Earth Sciences, v. 91, p. 1017–1029, doi: 10.1007/s00531-002-0294-7. Kryza, R., Mazur, S., and Pin, C., 1995, Leszczyniec meta-igneous complex in the eastern part of the Karkonosze Izera Block, western Sudetes: Trace element and Nd isotope study: Neues Jahrbuch für Mineralogie Abhandlungen, v. 169, p. 59–74. Kryza, R., Mazur, S., and Oberc-Dziedzic, T., 2004, The Sudetic geological mosaic: Insights into the root of the Variscan orogen: Przegla˛d Geologiczny, v. 52, p. 761–773. Lange, U., Bröcker, M., Mezger, K., and Don, J., 2002, Geochemistry and RbSr geochronology of a ductile shear zone in the Orlica-S´niez˙nik dome
(west Sudetes, Poland): International Journal of Earth Sciences, v. 91, p. 1005–1016, doi: 10.1007/s00531-002-0281-z. Lange, U., Bröcker, M., Armstrong, R., Z˙ elaz´niewicz, A., Trapp, E., and Mezger, K., 2005, The orthogneisses of the Orlica-S´niez˙nik complex (west Sudetes, Poland): Geochemical characteristics, the importance of pre-Variscan migmatization and constraints on the cooling history: Journal of the Geological Society of London, v. 162, p. 973–984, doi: 10.1144/0016-764904-095. Linnemann, U., Gehmlich, M., Tichomirowa, M., Buschmann, B., Nasdala, L., Jonas, P., Lützner, H., and Bombach, K., 2000, From Cadomian subduction to Early Palaeozoic rifting: The evolution of Saxo-Thuringia at the margin of Gondwana in the light of single zircon geochronology and basin development (central European Variscides, Germany), in Franke, W., Haak, V., Oncken, O., and Tanner, D., eds., Orogenic processes: Quantification and modelling in the Variscan Belt: London, Geological Society of London Special Publication 179, p. 131–153. Mazur, S., 1995, Structural and metamorphic evolution of the country rocks at the eastern contact of the Karkonosze granite in the southern Rudawy Janowickie Mts. and Lasocki Range: Geologia Sudetica, v. 29, p. 31–98. Mazur, S., and Aleksandrowski, P., 2001, The Teplá(?)/Saxothuringian suture in the Karkonosze-Izera Massif, western Sudetes, central European Variscides: International Journal of Earth Sciences, v. 90, p. 341–360, doi: 10.1007/s005310000146. Mazur, S., and Kryza, R., 1996, Superimposed compressional and extensional tectonics in the Karkonosze-Izera Block, NE Bohemian Massif, in Oncken, O., and Janssen, C., eds., Basement tectonics, v. 11: Dordrecht, Kluwer Academic Publishers, p. 51–66. Miller, C.F., McDowell, S.M., and Mapes, R.W., 2003, Hot and cold granites? Implications of zircon saturation temperatures and preservation of inheritance: Geology, v. 31, p. 529–532. Mingram, B., Kröner, A., Hegner, E., and Krentz, O., 2004, Zircon ages, geochemistry, and Nd isotopic systematics of pre-Variscan orthogneisses from the Erzgebirge, Saxony (Germany), and geodynamic interpretation: International Journal of Earth Sciences, v. 93, p. 706–727, doi: 10.1007/ s00531-004-0414-7. Nakamura, N., 1974, Determination of REE, Ba, Fe, Mg, Na and K in carbonaceous and ordinary chondrites: Geochimica et Cosmochimica Acta, v. 38, p. 757–775, doi: 10.1016/0016-7037(74)90149-5. Nance, R.D., Murphy, J.B., Strachan, R.A., D’Lemos, R.S., and Taylor, G.K., 1991, Late Proterozoic tectonostratigraphic evolution of the Avalonian and Cadomian terranes: Precambrian Research, v. 53, p. 41–78, doi: 10.1016/0301-9268(91)90005-U. Oberc-Dziedzic, T., Klimas, K., Kryza, R., and Fanning, C.M., 2003a, SHRIMP zircon geochronology of the Strzelin gneiss, SW Poland: Neoproterozoic thermal event in the fore-Sudetic block, central European Variscides: International Journal of Earth Sciences, v. 92, p. 701–711, doi: 10.1007/ s00531-003-0345-8. Oberc-Dziedzic, T., Klimas, K., Kryza, R., Fanning, C.M., and Madej, S., 2003b, SHRIMP zircon ages from gneiss help locate the west-east Sudetes boundary (NE Bohemian Massif, SW Poland): Journal of the Czech Geological Society, v. 48, no. 1–2, p. 98. Oberc-Dziedzic, T., Pin, C., and Kryza, R., 2005a, Early Palaeozoic crustal melting in an extensional setting: Petrological and Sm-Nd evidence from the Izera granite-gneisses, Polish Sudetes: International Journal of Earth Sciences, v. 94, p. 354–368, doi: 10.1007/s00531-005-0507-y. Oberc-Dziedzic, T., Kryza, R., Klimas, K., Fanning, M.C., and Madej, S., 2005b, Gneiss protolith ages and tectonic boundaries in the Variscan collage of the NE part of the Bohemian Massif (fore-Sudetic block, SW Poland): Geological Quarterly, v. 49, no. 4, p. 363–378. Okrusch, M., Seidel, E., Schüssler, U., and Richter, P., 1989, Geochemical characteristics of metabasites in different tectonic units of the northeast Bavarian crystalline basement, in Emmermann, R., and Wohlenberg, J., eds., The German continental deep drilling program (KT): Berlin, Springer, p. 67–79. Oliver, G.J.H., Corfu, F., and Krogh, T.E., 1993, U-Pb ages from SW Poland: Evidence for a Caledonian suture zone between Baltica and Gondwana: Journal of the Geological Society of London, v. 150, p. 355–369. Patiño Douce, A.E., and Beard, J.S., 1995, Dehydration melting of biotite gneiss and quartz amphibolite from 3 to 15 kbar: Journal of Petrology, v. 36, p. 707–738. Patiño Douce, A.E., and Johnston, A.D., 1991, Phase equilibria and melt productivity in the pelitic system: Implications for the origin of peraluminous
Late Cambrian (ca. 500 Ma) felsic anorogenic magmatism in the Bohemian Massif granitoids and aluminous granulites: Contributions to Mineralogy and Petrology, v. 107, p. 202–218, doi: 10.1007/BF00310707. Patocˇka, F., Dostal, J., and Pin, C., 1997, Early Palaeozoic intracontinental rifting in the central west Sudetes, Bohemian Massif: Geochemical and Sr-Nd isotopic study on felsic-mafic metavolcanics of the Rychory Mts. complex: Terra Nova, v. 9, abstract supplement 1, p. 144–145. Pearce, J.A., Harris, N.B.W., and Tindle, A.G., 1984, Trace element discrimination diagrams for the tectonic interpretation of granitic rocks: Journal of Petrology, v. 25, p. 956–983. Petford, N., and Gallagher, K., 2001, Partial melting of mafc (amphibolitic) lower crust by periodic influx of basaltic magma: Earth and Planetary Science Letters, v. 193, p. 483–499, doi: 10.1016/S0012-821X(01)00481-2. Pin, C., 1990, Variscan oceans: Ages, origins and geodynamic implications inferred from geochemical and radiometric data: Tectonophysics, v. 177, p. 215–227, doi: 10.1016/0040-1951(90)90282-D. Pin, C., and Santos Zalduegui, J.F., 1997, Sequential separation of light rare earth elements, thorium and uranium by miniaturized extraction chromatography: Application to isotopic analyses of silicate rocks: Analytica Chimica Acta, v. 339, p. 79–89, doi: 10.1016/S0003-2670(96)00499-0. Pin, C., Mierzejewski, M., and Duthou, J.L., 1987, Wiek izochronowy RbSr granitu karkonoskiego z kamieniołomu Szklarska Poreba Huta oraz oznaczenie stosunku inicjalnego 87Sr-86Sr w tymz˙e granicie: Przegla˛d Geologiczny, v. 35, p. 512–517. Poitrasson, F., Duthou, J.L., and Pin, C., 1995, The relationship between petrology and Nd isotopes as evidence for contrasting anorogenic granite genesis: Example of the Corsican Province (SE France): Journal of Petrology, v. 36, p. 1251–1274. Rötzler, K., Schumacher, R., Maresch, V.W., and Willner, A.P., 1998, Contrasts in the metamorphic evolution and exhumation history of different metapelite units in the western Erzgebirge (Saxony, Germany): European Journal of Mineralogy, v. 10, p. 261–280. Sandiford, M., and Powell, R., 1986, Deep crustal metamorphism during continental extension: Modern and ancient examples: Earth and Planetary Science Letters, v. 79, p. 151–158, doi: 10.1016/0012-821X(86)90048-8. Schmädicke, E., and Evans, B.W., 1997, Garnet-bearing ultramafic rocks from the Erzgebirge, and their relation to other settings in the Bohemian Massif: Contributions to Mineralogy and Petrology, v. 127, p. 57–74, doi: 10.1007/s004100050265. Seston, R., Winchester, J.A., Piasecki, M.A., Crowley, Q.G., and Floyd, P.A., 2000, A structural model for the western-central Sudetes: A deformed stack of Variscan thrust sheets: Journal of the Geological Society of London, v. 157, p. 1155–1167. Siebel, W., Raschka, H., Irber, W., Kreuzer, H., Lenz, K.L., Höhndorf, A., and Wendte, I., 1997, Early Palaeozoic acid magmatism in the Saxothuringian Belt: New insights from a geochemical and isotopic study of orthogneisses and metavolcanic rocks from the Fichtelgebirge, SE Germany: Journal of Petrology, v. 38, p. 203–230. Skjerlie, K.P., and Johnston, A.D., 1992, Vapour-absent melting at 10 kbar of a biotite-and amphibole-bearing tonalitic gneiss: Implications for the generation of A-type granites: Geology, v. 20, p. 263–266, doi: 10.1130/00917613(1992)020<0263:VAMAKO>2.3.CO;2. Skjerlie, K.P., and Patiño Douce, A.E., 1995, Anatexis of interlayered amphibolite and pelite at 10 kbar: Effect of diffusion of major components on phase relations and melt fraction: Contributions to Mineralogy and Petrology, v. 122, p. 62–78, doi: 10.1007/s004100050113. Smith, A.D., and Lewis, C., 1999, The planet beyond the plume hypothesis: Earth Science Reviews, v. 48, p. 135–182, doi: 10.1016/S00128252(99)00049-5. Štipská, P., Schulmann, K., and Kröner, A., 2004, Vertical extrusion and middle crustal spreading of omphacite granulite: A model of syn-convergent exhumation (Bohemian Massif, Czech Republic): Journal of Metamorphic Geology, v. 22, p. 179–198, doi: 10.1111/j.1525-1314.2004.00508.x. 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, doi: 10.1016/S0009-2541(00)00198-4. Taylor, S.R., and McLennan, S.M., 1985, The continental crust: Its composition and evolution: Oxford, Blackwell Scientific, 312 p. Thompson, A.B., 1996, Fertility of crustal rocks during anatexis: Transactions of the Royal Society of Edinburgh: Earth Sciences, v. 87, p. 1–10.
229
Thompson, A.B., 2000, Some space-time relationships for crustal melting and granitic intrusion at various depths, in Castro, A., Fernandez, J.L., and Vigneresse, J.L., eds., Understanding granites: Integrating new and classical techniques: London, Geological Society of London Special Publication 168, p. 7–25. Thompson, A.B., and Connolly, J.A.D., 1995, Melting of the continental crust: Some thermal and petrological constraints on anatexis in continental collision zones and other tectonic settings: Journal of Geophysical Research, v. 100, p. 15,565–15,579. Thompson, R.N., 1982, British Tertiary volcanic province: Scottish Journal of Geology, v. 18, p. 49–107. Tichomirowa, M., Berger, H.-J., Koch, E.A., Belyatski, B.V., Götze, J., Kempe, U., Nasdala, L., and Schaltegger, U., 2001, Zircon ages of high-grade gneisses in the eastern Erzgebirge (central European Variscides)—Constraints on origin of the rocks and Precambrian to Ordovician magmatic events in the Variscan foldbelt: Lithos, v. 56, p. 303–332, doi: 10.1016/S0024-4937(00)00066-9. Timmermann, H., Parrish, R.H., Noble, S.R., Kryza, R., and Patocˇka, F., 1999, Single cycle Variscan orogeny inferred from new U-(Th-)Pb data from the Sudetes mountains in Poland and the Czech Republic, in Abstracts of the PACE mid-term review and 4th PACE network meeting, October, Copenhagen: Copenhagen, Geological Institute, University of Copenhagen, v. 9–10, p. 24. Turniak, K., Mazur, S., and Wysoczanski, R., 2000, SHRIMP zircon geochronology and geochemistry of the Orlica-S´niez˙nik gneisses (Variscan Belt of central Europe) and their tectonic implications: Geodinamica Acta, v. 13, p. 293–312, doi: 10.1016/S0985-3111(00)01045-7. Vidal, P., Auvray, B., Charlot, R., Fediuk, F., Hameurt, J., and Waldhausrová, J., 1975, Radiometric age of volcanics of the Cambrian “Krˇivoklát Rokycany” complex (Bohemian Massif): Geologische Rundschau, v. 64, p. 563–570, doi: 10.1007/BF01820683. Waldhausrová, J., 1971, The chemistry of the Cambrian volcanics in the Barrandian area: Krystalinikum, v. 8, p. 45–75. Whalen, J.B., Currie, K.L., and Chappell, B.W., 1987, A-type granites: Geochemical characteristics, discrimination and petrogenesis: Contributions to Mineralogy and Petrology, v. 95, p. 407–419, doi: 10.1007/BF00402202. Willner, A.P., Rötzler, K., and Maresch, W.V., 1997, Pressure-temperature and fluid evolution of quartzo-feldspathic metamorphic rocks with a relic high-pressure granulite-facies history from the central Erzgebirge (Saxony, Germany): Journal of Petrology, v. 38, p. 307–336, doi: 10.1093/ petrology/38.3.307. Winchester, J.A., Floyd, P.A., Awdankiewicz, M., Piasecki, M.A.J., Awdankiewicz, H., Gunia, P., and Gliwicz, T., 1998, Geochemistry and tectonic significance of metabasic suites in the Góry Sowie block, SW Poland: Journal of the Geological Society of London, v. 155, p. 155–164. Z˙ elaz´niewicz, A., 1990, Deformation and metamorphism in the Góry Sowie gneiss complex, Sudetes, SW Poland: Neues Jahrbuch für Mineralogie Abhandlungen, v. 179, p. 129–157. Z˙ elaz´niewicz, A., 1994, Ductile shear zones versus amphibolitized basic dikes within the Izera gneisses: Fragment of a passive margin stretching from Lusatian to west Sudetes: Terra Nostra, v. 3, no. 94, p. 108–110. Z˙ elaz´niewicz, A., Nowak, I., Achramowicz, S., and Czaplin´ski, W., 2003, The northern part of the Izera-Karkonosze block: A passive margin of the Saxothuringian terrane, in Cie˛z˙kowski, W., Wojewoda, J., and Z˙ elaz´niewicz, A., eds., Sudety Zachodnie: od wendu do czwartorze˛du: Wrocław, WIND, p. 17–32. Z˙ elaz´niewicz, A., Dörr, W., Bylina, P., Franke, W., Haack, U., Heinisch, H., Schastok, J., Grandmontagne, K., and Kulicki, C., 2004, The eastern continuation of the Cadomian orogen: U–Pb zircon evidence from SaxoThuringian granitoids in south-western Poland and the northern Czech Republic: International Journal of Earth Sciences, v. 93, p. 773–781, doi: 10.1007/s00531-004-0418-3. Zulauf, G., Dörr, W., Fiala, J., and Vejnar, Z., 1997, Late Cadomian crustal tilting and Cambrian transtension in the Teplá-Barrandian unit (Bohemian Massif, central European Variscides): Geologische Rundschau, v. 86, p. 571–584. 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
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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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).
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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
TABLE 1. MAJOR AND TRACE ELEMENT DATA FOR THE BARRANDIAN METAVOLCANICS Depleted mantle source Mildly enriched mantle source Strongly enriched mantle source First subgroup Second subgroup Lhv-2 Rou-1 Kruš-1 Zb-1 Chr-1 Kli-1 La-2 Li-1 Reb-1 UP-1 Lit-1 Lit-2 UP-2 Zchl-1 Kli-3 Kot-1 Boro-1 Mitov-1 Mitov-2 Si-2
Viš-1
21.3 33.5 5.34 21.9 4.75 1.09 4.92 0.84 5.68 1.28 3.87 0.60 4.01 0.64 47.3
20 341 178 517
43 53.9 1905
20.6 43.5 5.30 21.1 4.55 1.28 4.15 0.64 3.80 0.74 2.13 0.33 2.20 0.35 20.5
17 104 37 126
49 119 450
21.0 43.7 5.25 20.9 4.40 1.16 4.05 0.63 3.75 0.75 2.14 0.33 2.15 0.34 21.3
16 104 39 121
51 154 768
63.7 63.7 57.1 10.3 15.0 14.0 5.81 6.50 6.18 0.06 0.085 0.124 1.38 3.05 2.82 1.61 1.77 6.33 2.29 4.07 3.21 1.68 1.61 1.66 0.41 0.87 0.81 0.05 0.17 0.16 12.8 3.78 7.60 100.08 100.65 99.94
Kruš-2 Nml-1
Sedimentary rocks
Zr 42.5 79.8 78.8 49.9 36.7 68.1 107 52.7 56.2 115 77.7 66.1 88.3 91.8 68.6 114 296 81.0 77.0 260 272 228 150 152 149 Hf 1.38 2.22 2.17 1.42 1.14 1.85 2.89 1.55 1.70 3.23 2.15 1.78 2.40 2.36 1.79 2.51 6.55 2.11 2.06 6.22 6.38 5.59 3.95 4.19 4.04 Nb 0.37 2.50 1.89 0.91 0.90 1.14 1.87 0.69 1.65 1.72 0.95 1.64 7.17 6.61 3.97 6.25 39.8 19.5 18.7 66.0 68.0 8.19 7.73 8.06 7.68 Ta 0.04 0.228 0.165 0.082 0.075 0.09 0.16 0.051 0.14 0.148 0.091 0.224 0.58 0.529 0.497 0.48 2.98 1.46 1.45 4.91 4.94 0.707 0.73 0.70 0.67 Th 0.042 0.23 0.175 0.085 0.107 0.108 0.23 0.098 0.225 0.20 0.132 0.35 0.65 0.58 0.55 0.59 3.80 1.8 1.79 6.46 6.64 7.63 6.24 4.74 5.01 U 0.516 0.484 0.096 0.512 0.163 0.068 1.314 7.411 0.507 1.244 0.653 0.189 0.239 0.398 0.463 0.173 1.522 0.442 0.439 1.517 1.592 2.438 10.69 1.69 1.95 Note: Major elements are in oxide wt%; trace elements in ppm. Measurements are done by inductively coupled plasma (ICP) atomic emission spectrometry for the major elements and ICP mass spectrometry for the trace elements, at the Centre de Recherche Pétrographique et Géochimique, Nancy, France, following methods described by Carignan et al. (2001). L.D.—limit of detection; LOI—loss on ignition; nd—not determined.
4.18 12.2 1.96 10.5 3.60 1.22 4.77 0.83 5.43 1.13 3.21 0.48 3.09 0.46 31.6
37 138 53 282
2 150 137
3 75.2 74
(ppm) Rb Sr Ba
Teb-1
46.5 53.7 16.00 13.4 9.97 9.2 0.2 0.11 7.68 6.94 8.05 8.19 2.04 4.37 0.11 0.16 0.91 1.36 0.11 0.18 8.49 2.24 100.07 99.83
Kli-5
(wt%) SiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O TiO2 P2O5 LOI Total
Sample
Sm-Nd isotope and trace element study of Late Proterozoic metabasalts 235
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TABLE 2. Sm AND Nd CONCENTRATIONS 147 144 143 144 Sm/ Nd Nd/ Nd Sample Sm Nd εNd570Ma εNd600Ma Depleted mantle source First subgroup Kli-5 2.00 4.67 0.2587 0.513328 (7) +8.9 +8.7 Teb-1 3.79 11.2 0.2052 0.513081 (6) +8.0 +8.0 Lhv-2 3.75 10.8 0.2096 0.513134 (6) +8.7 +8.7 Rou-1 1.96 5.54 0.2137 0.513165 (11) +9.0 +8.9 Kruš-1 1.67 4.11 0.2448 0.513305 (9) +9.5 +9.3 Zb-1 2.73 7.22 0.2285 0.513212 (8) +8.8 +8.7 Second subgroup Chr-1 4.38 12.3 0.2147 0.513112 (6) +7.9 +7.8 Kli-1 2.32 5.50 0.2544 0.513308 (7) +8.8 +8.6 La-2 2.61 7.13 0.2215 0.513183 (7) +9.0 +8.7 Li-1 4.71 12.8 0.2217 0.513145 (6) +8.0 +7.9 Reb-1 3.15 8.74 0.2178 0.513171 (7) +8.8 +8.7 UP-1 2.56 7.34 0.2111 0.513097 (6) +7.9 +7.8 Mildly enriched mantle source Lit-1 3.23 11.4 0.1711 0.512883 (8) +6.6 +6.7 Lit-2 3.37 11.9 0.1717 0.512938 (7) +7.6 +7.7 UP-2 2.50 8.09 0.1869 0.513005 (6) +7.8 +7.9 Zchl-1 3.48 12.2 0.1724 0.512964 (8) +8.1 +8.2 Kli-3 7.88 38.2 0.1246 0.512551 (7) +3.5 +3.8 Strongly enriched mantle source Kot-1 3.24 14.4 0.1362 0.512666 (9) +4.9 +5.1 Boro-1 3.56 15.8 0.1358 0.512653 (7) +4.7 +4.8 Mit-1 8.78 43.4 0.1224 0.512568 (7) +4.0 +4.3 Mit-2 8.76 43.3 0.1223 0.512573 (8) +4.1 +4.4 Si-2 4.69 23.0 0.1232 0.512513 (8) +2.9 +3.2 Sedimentary rocks Viš-1 4.65 22.8 0.1235 0.512318 (8) –1.0 –0.7 Nml-1 4.76 22.4 0.1235 0.512369 (7) 0.0 –0.1 Kruš-2 5.15 24.0 0.1295 0.512225 (8) –3.2 –3.0 Note: εNd values are calculated based on the chondritic parameters of Jacobsen and Wasserburg (1980). Concentrations are in μg/g and are measured by isotope dilution and thermal ionization mass 147 144 spectrometry. Sm/ Nd ratios are precise to 0.2% at the 95% 143 144 confidence level. Nd/ Nd ratios are corrected for mass-dependent 146 144 fractionation by normalization to Nd/ Nd = 0.7219. The quoted precision is based on within-run statistics and corresponds to the standard error on the mean at the 95% confidence level. During the period of 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 143 144 Nd/ Nd = 0.511857 for the widely used La Jolla isotopic standard (Tanaka et al., 2000). See Appendix for sample locations.
considered to be essentially immobile during zeolite- to greenschist-facies transformations (e.g., Wood et al., 1979; Saunders et al., 1980) have been considered in this work. These include high-field-strength elements (HFSE: Ti, Zr, Nb; e.g., Cann, 1970; Dungan et al., 1983); Th (Saunders and Tarney, 1984; Chen et al., 1986; Krauskopf, 1986); the rare earth elements (REE; Herrmann et al., 1974; Menzies et al., 1977, 1979; Dungan et al., 1983); and transition elements (Cr, Ni), including V (Shervais, 1982). However, significant enrichment of REE (particularly heavy REE [HREE]) and Th has been shown to have occurred in some extremely altered samples in which the HFSE remained essentially undisturbed (e.g., Wharton et al., 1995). For this reason, only least-altered material was selected for this study. Sm and Nd, adjacent elements of the chemically coherent lanthanide
group, are involved in a radiogenic isotope system as parent and daughter isotopes, respectively, and so display a very high degree of chemical congruence. This congruence implies a very reduced relative mobility, thereby favoring a great resistance of the Sm-Nd isotopic pair with regard to geological disturbances (e.g., DePaolo, 1988). Indeed, it has been shown that seawater alteration is not able to fractionate Sm from Nd and disturb Nd isotope systematics (e.g., Staudigel and Hart, 1983), unless extremely high (on the order of 105) fluid/rock ratios are achieved (e.g., Jacobsen and Wasserburg, 1979). These conditions are presumably restricted to the immediate surroundings of hydrothermal vents and are not relevant to the samples investigated in this work. Based on these published lines of evidence, it is inferred that the chemical and Nd isotope features discussed in the following sections may be ascribed to the primary, igneous history of the studied rocks. Immobile, incompatible trace element data are shown as chondrite-normalized patterns, in order to depict their relative fractionation and highlight any “anomalous” behavior of certain diagnostic elements. Besides REE, these diagrams include the even more incompatible elements Th and Nb, which allows the degree of depletion of mantle sources to be assessed. In addition, the relative fractionation of Th, Nb, and light REE (LREE) is of great petrogenetic significance because it may provide evidence (as negative anomalies of Nb) for the involvement of a subduction component during the petrogenesis of ancient, altered basaltic rocks (e.g., Jenner et al., 1991). Indeed, supra-subduction zone lavas are almost invariably characterized by relatively high Th/Nb ratios (e.g., Sun, 1980) compared to mid-ocean ridge basalts (MORB; both of normal and enriched types) and withinplate basalts, which display nearly constant and very low values (~0.05–0.08; e.g., Sun, 1980). Alternatively, elevated Th/Nb ratios in basaltic magmas may reflect merely contamination by wall-rock assimilation during magma ascent and emplacement (e.g., Dupuy and Dostal, 1984) and/or the addition of materials derived from the continental crust, because continental materials are characterized by a strong enrichment of Th and LREE relative to Nb (e.g., Taylor and McLennan, 1985). Nd isotopes provide a useful tool to assist the choice to be made between these alternative interpretations (e.g., Pin and Paquette, 1997), insofar as isotopically distinctive, old sialic crust was involved. Several subgroups can be distinguished among the Barrandian metabasalts and further subdivided on the basis of their chondrite-normalized trace element patterns: 1. One subgroup gathers six samples (Klí-5, Teb-1, Lhv-2, Rou-1, Kruš-1, and Zb-1; Fig. 2A) with almost flat patterns (at ~10× chondrite abundances) for moderately incompatible elements (specifically, elements from Nd to Lu, ranked as a function of increasing bulk partition coefficients during partial melting of upper mantle mineral assemblages). This subgroup has a strong relative depletion of the most incompatible elements (Th, Nb, La, Ce). In detail, the heavy REE patterns range from flat (Klí-5, Kruš-1) to slightly fractionated (Teb-1, Rou-1).
Sm-Nd isotope and trace element study of Late Proterozoic metabasalts 100
Rock/chondrites
Rock/chondrites
100
10
1
10
1
B
A 0.1
237
Th Nb La Ce Nd Zr Sm Eu Ti Gd Dy Y Er Yb Lu
0.1
Th Nb La Ce Nd Zr Sm Eu Ti Gd Dy Y Er Yb Lu
Figure 2. Chondrite-normalized patterns of selected immobile incompatible trace elements (A) for the first subgroup of depleted metabasalts, without significant negative Nb anomaly and (B) for the second subgroup of depleted metabasalts, showing a weak negative anomaly of Nb. See text for further comments. Normalization values from Sun and McDonough (1989).
All these samples display a distinct negative anomaly for Ti, combined in some cases (Kruš-1, Klí-5, Teb-1) with a faint negative anomaly of Zr. There is no significant Nb anomaly in these samples. Zr/Nb ratios range from 32 to 60 with an extremely high value (115) in sample Kli-5 (mean 58, standard deviation [SD] 30). These elevated ratios point to a strongly depleted mantle source (e.g., Le Roex et al., 1983). Overall, this group of samples, characterized by a strong depletion of the most incompatible trace elements, shows clear affinities with normal MORB (N-MORB), although the ubiquitous Ti anomalies, occasional Zr negative anomalies, and relatively low Ti/V ratios (range 20–30, mean 25, SD 4) are reminiscent of some supra-subduction zone influence. Such a subtle influence is corroborated by Th/Nb ratios (range 0.09–0.12, mean 0.10, SD 0.01) slightly higher than the value typical for N-MORB, ocean island basalts (OIB), and transitional magma types (~0.07). Initial Nd isotope signatures (εNd600) calculated for these samples are all extremely radiogenic (from +8.0 to +9.3, mean +8.7, SD 0.4) and demonstrate that the mantle source was strongly depleted in Nd relative to Sm on a secular basis. In this respect, the mafic melts parental to the Barrandian “spilites” might have been extracted from a typical N-MORB mantle source, which was characterized, 600 Ma ago, by εNd values of approximately +9, according to the evolutionary model of Zindler (1982). In any case, the basaltic magmas did not suffer any significant interaction with ancient crustal materials characterized by unradiogenic Nd isotope composition (i.e., distinctly negative εNd values). 2. A second subgroup of six samples (Chrˇ-1, Klí-1, La-2, Li-1, Reb-1, and UP-1) displays the same trace element characteristics, that is, flat chondrite-normalized patterns
from Nd to Lu at ~10 times the chondritic level, with negative Ti anomalies and strong relative depletion of the most incompatible elements from Th to Ce (Fig. 2B). However, relative abundances of Th tend to be higher, as shown by Th/Nb ratios ranging from 0.12 to 0.21 (mean 0.15, SD 0.03) and illustrated by faint negative Nb anomalies on chondrite-normalized diagrams. Otherwise, these samples share the same characteristics as the previous group, that is, elevated Zr/Nb ratios (34–82, mean 60, SD 20), moderate Ti/V ratios (19–27, mean 23, SD 3), and very radiogenic values (from +7.8 to +8.7, mean 8.3, SD 0.4). The sample with the deepest Nb anomaly (UP-1) also has a somewhat lower εNd value (+7.8), possibly suggesting that this sample suffered some crustal contamination by materials with high Th/Nb and low εNd, such as the graywackes forming most of the Barrandian Proterozoic (e.g., samples Nml-1 and Vis-1, with Th/Nb ~0.60 and εNd600 approximately −1). However, sample Chrˇ-1, with an identical εNd600 of + 7.8, does not display an elevated Th/Nb ratio (0.12), suggesting that mantle source heterogeneity might provide an alternative explanation for the observed range of εNd values. Based on these data, it is apparent that there is no clear compositional gap between these two subgroups of samples. Indeed, the broad continuity of this set of samples is highlighted by multielement patterns normalized to average N-MORB. In this diagram (Fig. 3), the samples display almost flat patterns from Nd to Lu, with a general relative depletion of Zr and Ti. For the most incompatible elements, their patterns continuously range from strongly LREE-depleted to very slightly LREE-enriched, but all show a marked negative Nb anomaly relative to La, and especially to Th. Such chemical features demonstrate that, although similar to N-MORB in many respects, these metabasalts are characterized by a relative depletion in HFSE, particularly Nb. For this
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Pin and Waldhausrová 100
Rock/chondrites
Rock/NMORB
10
1
0.1
Th Nb La Ce Nd Zr Sm Eu Ti Dy Y Yb Lu
Figure 3. N-MORB–normalized patterns of selected immobile incompatible trace elements for the samples of depleted metabasalts. Note the large negative anomaly of Nb throughout, and the LREE fractionation in the most-depleted samples, interpreted to reflect derivation from a strongly depleted mantle source that has been fluxed by a component (hydrous silicic fluid and/or melt?) rich in Th and poor in Nb. Normalization values from Sun and McDonough (1989).
reason, these samples are hereafter collectively referred to as “Nb-depleted basalts,” or simply, “depleted basalts.” Although this term is admittedly of limited significance, note that eight of the twelve samples in this group (i.e., Lhv-2, Zb-1, Klí-1, Klí-5, La-2, Reb-1, Rou-1, and Kruš-1) plot along a rather poor isochron (mean square of weighted deviations [MSWD] = 3.7) in a 143Nd/144Nd versus 147Sm/144Nd diagram (not shown). This isochron indicates an age of 605 ± 39 Ma (uncertainty quoted at the 95% confidence level, using the Isoplot software; Ludwig, 1994), with an initial εNd of +8.8 ± 0.2, as calculated following the method of Fletcher and Rossman (1982). The four remaining samples (Chrˇ -1, Li-1, Teb-1, and UP-1) plot along a linear array (MSWD = 0.5) with a similar slope and a lower initial ratio, but no meaningful age (598 ± 162 Ma) can be obtained because of the limited number of samples and the small spread of Sm/Nd ratios. Nevertheless, in the absence of better radiometric evidence, the imprecise date of 605 ± 39 Ma is considered to reflect the igneous emplacement of the depleted metabasalts, in broad agreement with the Ediacaran age suggested by micropaleontological data for associated sediments. 3. A third subgroup comprising four samples (Lit-1, Lit-2, Zchl-1, and to some extent, UP-2) clearly depart from the previous rocks in showing chondrite-normalized REE patterns gently sloping from La to Lu, that is, a fractionation pattern opposite to that observed in the previous samples. Although negative Ti anomalies occur throughout (Fig. 4), the other HFSE abundances are in marked contrast with those of the previous samples. Indeed, Nb is much more abundant in these rocks and does not show any depletion, but instead an enrichment relative to adjacent elements on chondrite-normalized diagrams (Fig. 4), whereas Th
10
1
Th Nb La Ce Nd Zr Sm Eu Ti Gd Dy Y Er Yb Lu
Figure 4. Chondrite-normalized patterns of selected immobile incompatible trace elements for metabasalts mildly enriched in LREE (see text). Normalization values from Sun and McDonough (1989).
contents are almost constant at ~10 times chondrite. Zr/ Nb ratios are much lower (12–18) than those measured in the previous, depleted, group and clearly suggest a different mantle source. Th/Nb ratios show very little scatter (0.09–0.10) and preclude significant degrees of crustal assimilation or recycling into the mantle source. Overall, the REE and HFSE features indicate derivation from a mildly enriched mantle. However, εNd600 values are still strongly positive (from +6.7 to +8.2), implying that the mantle source of these rocks was depleted in Nd relative to Sm on a time-integrated basis. 4. The last group of samples comprises five metabasalts (Boro-1, Kli-3, Kot-1, Mit-1, and Mit-2) and one felsic lava (“keratophyre,” Si-2) classified as a trachyte according to the TAS diagram (Le Maitre et al., 1989). All the metabasalts display strong enrichment in LREE, Th, and especially Nb, which gives rise to positive anomalies on chondrite-normalized diagrams (Fig. 5A). Along with very low Zr/Nb ratios (4–7) and, with two exceptions, high Ti/V (51–77), these geochemical features resemble those of alkaline basalts, extracted from strongly enriched mantle sources. Two samples (Boro-1 and Kot-1) show negative anomalies of Ti and Zr, associated with much lower Ti/V (25, 27) and interpreted to reflect fractionation of Fe-Ti oxide rather than being intrinsic to parental magmas. Indeed, the felsic derivative (Si-2, with 70 wt% SiO2) ascribed to the same group of samples has deep anomalies of Ti and Nb (and of Sr and Eu), suggesting that fractionation of a plagioclase and Fe-Ti oxide (plus pyroxene and/or amphibole) assemblage was involved (Fig. 5B). As in the previous group of samples, low values of Th/Nb (0.09–0.10) put strict limitations on the involvement, if any, of crustal assimilation or source contamination processes. εNd600 values measured in the mafic samples range
Sm-Nd isotope and trace element study of Late Proterozoic metabasalts 1000
Rock/chondrites
Rock/chondrites
1000
100
10
100
10
B
A 1
239
1
Th Nb La Ce Nd Zr Sm Eu Ti Gd Dy Y Er Yb Lu
Th Nb La Ce Nd Zr Sm Eu Ti Gd Dy Y Er Yb Lu
Figure 5. Chondrite-normalized patterns of selected, immobile incompatible trace elements for metabasalts enriched in LREE and Nb (A) with clear affinity with OIBs and (B) for an associated metatrachyte Si-2. Normalization values from Sun and McDonough (1989).
1000
Rock/chondrites
from +3.8 to +5.1, whereas the metatrachyte has a slightly lower value (+ 3.2), which might allow for a subordinate amount of crustal assimilation during magma differentiation. Albeit significantly less radiogenic than those observed in the mildly enriched group of samples, the Nd isotope signatures of these alkaline basalts still reflect the time-integrated depletion of their mantle source. The number of samples analyzed is far too limited to allow any sound characterization of metasedimentary rocks. Instead we simply note that the two samples of graywackes (Nml-1 and Viš-1, collected from two localities separated by a few hundred meters) have essentially identical incompatible trace element abundances (specifically, large Th and LREE enrichment, unfractionated HREE, large negative anomalies of Nb and Ti, and very small negative Eu anomaly; Fig. 6). Sm-Nd isotope characteristics, with εNd600 values close to zero, preclude major derivation from old continental crust sources but instead require a significant input of juvenile (volcaniclastic?) material, in agreement with the immature characteristics (e.g., Na2O > 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.
REFERENCES CITED Anderson, D.L., 1982, Isotopic evolution of the mantle: The role of magma mixing: Earth and Planetary Science Letters, v. 57, p. 1–12, doi: 10.1016/0012821X(82)90168-6. Anderson, D.L., 1995, Lithosphere, asthenosphere, and perisphere: Reviews of Geophysics, v. 33, p. 125–149, doi: 10.1029/94RG02785.
Sm-Nd isotope and trace element study of Late Proterozoic metabasalts Bandres, A., Eguiluz, 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. Bandres, A., Eguiluz, L., Pin, C., Paquette, J.L., Ordoñez, B., Le Fèvre, B., Ortega, L.A., and Gil Ibarguchi, J.I., 2004, The northern Ossa-Morena Cadomian batholith (Iberian Massif): Magmatic arc origin and early evolution: International Journal of Earth Sciences, v. 93, p. 860–885, doi: 10.1007/s00531-004-0423-6. Bernardová, E., and Chab, J., 1974, Prehnite and pumpellyite in Upper Proterozoic basalts of the north-western part of the Barrandian region (Czechoslovakia): Krystalinikum, v. 10, p. 53–66. Bowring, S., Myrow, P., Landing, E., Ramezni, J., and Grotzinger, J., 2003, Geochronological constraints on terminal Neoproterozoic events and the rise of metazoans: Geophysical Research Abstracts, v. 5, p. 13,219. Cabanis B., and Wyns, R., 1986, Le volcanisme précambrien des Mauges (SE du Massif armoricain) et ses caractères géochimiques: Hercynica II, v. 1, p. 71–78. Cabanis, B., Chantraine, J., and Rabu, D., 1987, Geochemical study of the Brioverian (late Proterozoic) volcanic rocks in the northern Armorican Massif (France). Implications for geodynamic evolution during the Cadomian, in Pharaoh, T.C, Beckinsale, R.D., and Rickard D., eds., Geochemistry of mineralization of Proterozoic volcanic suites: London, Geological Society of London Special Publication 33, p. 503–523. Cann, J.R., 1970, Rb, Sr, Y, Zr and Nb in some ocean floor basaltic rocks: Earth and Planetary Science Letters, v. 10, p. 7–11, doi: 10.1016/0012821X(70)90058-0. Carignan, J., Hild, P., Mevelle, G., Morel, J., and Yeghicheyan, D., 2001, Routine analyses of trace elements in geological samples using flow injection and low-pressure on-line liquid chromatography coupled to ICP-MS: A study of geochemical reference materials BR, DR-N, UB-N, AN-G and GH: Geostandards Newsletter, v. 25, p. 187–198. Chaloupsky, J., Chlupacˇ, I., Mašek, J., Waldhausrová, J., and Chab, J., 1995, Teplá-Barrandian zone (Bohemicum). Stratigraphy, in Dallmeyer, R.D., Franke, W., and Weber, K., eds., Pre-Permian geology of central and eastern Europe: Berlin and Heidelberg, Springer-Verlag, p. 379–391. Chantraine, J., Egal, E., Thiéblemont, D., Le Goff, E., Guerrot, C., Ballèvre, M., and Guennoc, P., 2001, The Cadomian active margin (north Armorican Massif, France): A segment of the North Atlantic Panafrican belt: Tectonophysics, v. 331, p. 1–18, doi: 10.1016/S0040-1951(00)00233-X. Chen, J.H., Wasserburg, G.J., von Damm, K.L., and Edmond, J.M., 1986, The U-Th-Pb systematics in hot springs on the east Pacific rise at 21°N and Guaymas basin: Geochimica et Cosmochimica Acta, v. 50, p. 2467–2479, doi: 10.1016/0016-7037(86)90030-X. Cocherie, A., Chantraine, J., Fanning, C.M., Dabard, M.P., Paris, F., Le Herissé, and Egal, E., 2001, Datation U/Pb: âge Brivérien de la série d’Erquy (Massif armoricain, France). Compte-Rendus à l’Académie des Sciences de Paris, v. 333, p. 427–434. Cox, P.A., 1995, The elements on Earth: Oxford, Oxford University Press, 293 p. Dabard, M.P., 1990, Lower Brioverian formations (Upper Proterozoic) of the Armorican Massif (France): Geodynamic evolution of source areas revealed by sandstone petrography and geochemistry: Sedimentary Geology, v. 69, p. 45–58, doi: 10.1016/0037-0738(90)90100-8. Dabard, M.P., 2000, Petrogenesis of graphitic cherts in the Armorican segment of the Cadomian orogenic belt (NW France): Sedimentology, v. 47, p. 787–800, doi: 10.1046/j.1365-3091.2000.00316.x. Dabard, M.P., Loi, A., and Peucat, J., 1996, Zircon typology combined with Sm-Nd whole-rock isotope analysis to study Brioverian sediments from the Armorican Massif: Sedimentary Geology, v. 101, p. 243–260, doi: 10.1016/0037-0738(95)00068-2. DePaolo, D.J., 1988, Neodymium isotope geochemistry. An introduction: Berlin, Springer-Verlag, 187 p. Dörr, W., Zulauf, G., Fiala, J., Franke, W., and Vejnar, Z., 2002, Neoproterozoic to Early Cambrian history of an active plate margin in the Teplá Barradian unit—A correlation of U-Pb isotopic dilution-TIMS ages: Tectonophysics, v. 352, p. 65–85. Drost, K., Linnemann, U., McNaughton, N., Fatka, O., Kraft, P., Gehmlich, M., Tonk, C., and Marek, J., 2004, New data on the Neoproterozoic-Cambrian geotectonic setting of the Tepla-Barrandian volcano-sedimentary successions: Geochemistry, U-Pb zircon ages, and provenance (Bohemian Massif, Czech Republic): International Journal of Earth Sciences, v. 93, p. 742–757, doi: 10.1007/s00531-004-0416-5.
245
Dungan, M.A., Vance, J.A., and Blanchard, D.P., 1983, Geochemistry of the Shuksan greenschists and blueschists, North Cascades, Washington: Variably fractionated and altered metabasalts of oceanic affinity: Contributions to Mineralogy and Petrology, v. 82, p. 131–146, doi: 10.1007/ BF01166608. Dupret, L., Dissler, E., Doré, F., Gresselin, F., and Le Gall, J., 1990, Cadomian geodynamic evolution of the northeastern Armorican Massif (Normandy and Maine), in D’Lemos, R.S., Strachan, R.A., and Topley, C.G., eds., The Cadomian orogen: London, Geological Society of London Special Publication 51, p. 115–131. Dupuy, C., and Dostal, J., 1984, Trace element geochemistry of some continental tholeiites: Earth and Planetary Science Letters, v. 67, p. 61–69, doi: 10.1016/0012-821X(84)90038-4. Egal, E., Guerrot, C., Le Goff, E., Thiéblemont, D., and Chantraine, J., 1996, The Cadomian orogeny revisited in northern Brittany. Avalonian and related Peri-Gondwanan terranes of the Circum-North Atlantic: Boulder, Colorado, Geological Society of America Special Paper 304, p. 281–318. Eguiluz, L., Gil Ibarguchi, J.I., Abaloz, B., and Apraiz, A., 2000, Superposed Hercynian and Cadomian orogenic cycles in the Ossa-Morena zone and related areas of the Iberian Massif: Geological Society of America Bulletin, v. 112, p. 1398–1413, doi: 10.1130/0016-7606(2000)112<1398: SHACOC>2.0.CO;2. Eiler, J.M., McInnes, B., Valley, J.W., Graham, C.M., and Stolper, E.M., 1998, Oxygen isotope evidence for slab-derived fluids in the sub-arc mantle: Nature, v. 393, p. 777–781, doi: 10.1038/31679. Elsasser, W.M., 1971, Sea floor spreading and thermal convection: Journal of Geophysical Research, v. 76, p. 1101–1111. Fiala, F., 1977, The Upper Proterozoic volcanism of the Barrandian area and the problem of spilites: Sborník Geologických Veˇd, Geologie, v. 30, p. 2–247. Fiala, F., 1978, The TiO2-K2O-P2O5 diagram and tectonomagmatic relations of the volcanics of the Barrandian area: Veˇstnik Cˇeského geologického ustavu, v. 53, p. 333–346. Finger, F., and Pin, C., 1997, Arc-type crustal zoning in the Bruno-Vistulicum, eastern Czech Republic: A tale of the Late-Proterozoic Euro-Gondwana margin: Journal of the Czech Geological Society, v. 42, no. 3, p. 53. Finger, F., Hanzl, P., Pin, C., von Quadt, A., and Steyrer, H.P., 2000a, The Brunovistulian: Avalonian Precambrian at the eastern end of the central European Variscides? in Franke, W., Haak, V., Oncken, O., and Tanner, D., eds., Orogenic processes: Quantification and modelling in the Variscan belt: London, Geological Society of London Special Publication 179, p. 103–112. Finger, F., Tichomirowa, M., Pin, C., and Hanzl, P., 2000b, Relics of an early Pan-African metabasite-metarhyolite formation in the Brno Massif, Moravia, Czech Republic: International Journal of Earth Sciences, v. 89, p. 328–335, doi: 10.1007/s005310000084. Fletcher, I.R., and Rossman, K.J.R., 1982, Precise determination of εNd from Sm-Nd isochron: Geochemica et Cosmochimica Acta, v. 46, p. 1983– 1987, doi: 10.1016/0016-7037(82)90138-7. Fryer, P., Sinton, J., and Philpotts, J.A., 1981, Basaltic glasses from the Mariana Trough, in Hussong, D.M., and Uyeda, S., eds., Initial reports of the Deep Sea Drilling Project: Washington, D.C., National Science Foundation, v. 60, p. 601–610. Gill, J.B., 1984, Sr-Pb-Nd isotopic evidence that MORB and OIB sources contribute to ocean island arc magmas in Fiji: Earth and Planetary Science Letters, v. 68, p. 443–458, doi: 10.1016/0012-821X(84)90129-8. Gill, J.B., and Whelan, P., 1989, Postsubduction ocean island alkali basalts in Fiji: Journal of Geophysical Research, v. 94, p. 4579–4588. Glasmacher, U.A., Mann, U., and Wagner, G.A., 2002, Thermotectonic evolution of the Barrandian, Czech Republic, as revealed by apatite fissiontrack analysis: Tectonophysics, v. 359, p. 381–402, doi: 10.1016/S00401951(02)00538-3. Gradstein, F.M., Ogg, J.G., Smith, A.G., Bleeker, W., and Lourens, L.J., 2004, A new geologic time scale, with special reference to Precambrian and Neogene: Episodes, v. 27, p. 83–100. Gribble, R.F., Stern, R.J., Newman, S., Bloomer, S.H., and O’Hearn, T., 1998, Chemical and isotopic composition of lavas from the northern Mariana Trough: Implications for magma genesis in back-arc basins: Journal of Petrology, v. 39, p. 125–154, doi: 10.1093/petrology/39.1.125. Guerrot, C., and Peucat, J.-J., 1990, U-Pb geochronology of the Upper Proterozoic Cadomian orogeny in the northern Armorican Massif, France, in D’Lemos, R.S., Strachan, R.A., and Topley, C.G., eds., The Cadomian
246
Pin and Waldhausrová
orogen: London, Geological Society of London Special Publication 51, p. 13–26. Hamilton, W.B., 1995, Subduction systems and magmatism, in Smellie, J.L., ed., Volcanism associated with extension at consuming plate margins: London, Geological Society of London Special Publication 81, p. 3–28. Hawkins, J.W., 1995, Evolution of the Lau Basin—Insights from ODP Leg 135, in Taylor, B., and Natland, J., eds., Active margins and marginal basins of the Western Pacific: Washington, D.C., American Geophysical Union Geophysical Monograph 88, p. 125–173. Herrmann, A.G., Potts, M.J., and Knake, D., 1974, Geochemistry of the rare earth elements in spilites from the oceanic and continental crust: Contributions to Mineralogy and Petrology, v. 44, p. 1–16, doi: 10.1007/ BF00373128. Hochstaedter, A.G., Gill, J.B., and Morris, J.D., 1990, Volcanism in the Sumisu rift, II. Subduction and non-subduction related components: Earth and Planetary Science Letters, v. 100, p. 195–209, doi: 10.1016/0012821X(90)90185-Z. Hofmann, A.W., 1997, Mantle geochemistry: The message of oceanic volcanism: Nature, v. 385, p. 219–229, doi: 10.1038/385219a0. Holubec, J., 1995a, Structure of the Teplá-Barrandian zone (Bohemicum), in Dallmeyer, R.D., Franke, W., and Weber, K., eds., Pre-Permian geology of central and eastern Europe: Berlin and Heidelberg, Springer-Verlag, p. 392–397. Holubec, J., 1995b, Lithostratigraphic map of the Teplá-Barrandian Proterozoic. Internal report of the Geological Institute of the Czech Academy of Sciences. Jacobsen, S.B., and Wasserburg, G.J., 1979, Nd and Sr isotopic study of the Bay of Islands ophiolite complex and the evolution of the source of mid-ocean ridge basalts: Journal of Geophysical Research, v. 84, p. 7429–7445. Jacobsen, S.B., and Wasserburg, G.J., 1980, Sm-Nd isotopic evolution of chondrites: Earth and Planetary Science Letters, v. 50, p. 139–155, doi: 10.1016/0012-821X(80)90125-9. Jakeš, P., Zoubek, J., Zoubkova, J., and Franke, W., 1979, Graywackes and metagraywackes of the Teplá-Barrandian Proterozoic area: Sborník Geologických Veˇd, Geologie, v. 33, p. 83–122. Jenner, G., Dunning, G.R., Malpas, J., Brown, M., and Brace, T., 1991, Bay of Islands and Little Port complexes, revisited: Age, geochemical and isotopic evidence confirm supra-subduction zone origin: Canadian Journal of Earth Sciences, v. 28, p. 1635–1652. Kettner, R., 1918, Versuch einer Stratigraphischen Einteilung des Böhmischen Algonkiums: Geol. Rundschau, v. 8, p. 169–188, doi: 10.1007/BF01800895. Konzalová, M., 1980, Some Late Precambrian microfossils from the Bohemian Massif and their correlation: Precambrian Research. v. 15, p. 3–62. Konzalová, M., 2000, Organic-walled microbiota from the greywackes and other siliciclastic sediments of the Barrandian Neoproterozoic (Bohemian Massif, Czech Republic): Veˇstnik Cˇeského geologického ustavu, v. 75, p. 197–200. Krauskopf, K.B., 1986, Thorium and rare-earth metals as analogs for actinide elements: Chemical Geology, v. 55, p. 323–335, doi: 10.1016/00092541(86)90033-1. Krˇibek, B., Pouba, Z., Skocek, V., and Waldhausrová, J., 2000, Neoproterozoic of the Teplá-Barrandian unit as a part of the Cadomian orogenic belt: A review and correlation aspects: Veˇstnik Cˇeského geologického ustavu, v. 75, p. 175–196. Lees, G.J., Roach, R.A., Shufflebotham, M.M., and Griffiths, N.H., 1987, Upper Proterozoic basaltic volcanism in the northern Massif Armoricain, France, in Pharaoh, T.C, Beckinsale, R.D., and Rickard, D., eds., Geochemistry of mineralization of Proterozoic volcanic suites: London, Geological Society of London Special Publication 33, p. 503–523. Le Maitre, R.W., Bateman, P., Dudek, A., Keller, J., Lameyre, J., Le Bas, M.J., Sabine, P.A., Schmid, R., Sorensen, H., Treckeisen, A., Wooley, A.R., and Zanettin, B., 1989, A classification of igneous rocks and glossary of terms: Oxford, Blackwell Scientific, 193 p. Le Roex, A.P., Dick, H.J.B., Erlank, A.J., Reid, A.M., Fry, F.A., and Hart, S.R., 1983, Geochemistry, mineralogy, and petrogenesis of lavas erupted along the SW Indian ridge between the Bouvet triple junction and 11 degrees East: Journal of Petrology, v. 24, p. 267–318. Ludwig, K.R., 1994, Isoplot: A plotting and regression program for radiogenic isotope data, version 2.71: Reston, Virginia, U.S. Geological Survey Open File Report 91-445, 44 p. Marsaglia, K.M., Boggs, S., Jr., Clift, P., Seyedolali, A., and Smith, R., 1995, Sedimentation in western Pacific back-arc basins: New insights from
recent ODP drilling, in Taylor, B., and Natland, J., eds., Active margins and marginal basins of the western Pacific: Washington, D.C., American Geophysical Union Geophysical Monograph 88, p. 291–314. Martinez, F., and Taylor, B., 2002, Mantle wedge control on back-arc crustal accretion: Nature, v. 416, p. 417–442, doi: 10.1038/416417a. Mašek, J., 2000, Stratigraphy of the Proterozoic of the Barrandian area: Veˇstnik Cˇeského geologického ustavu, v. 75, p. 197–200. Menzies, M., Blanchard, D., and Jacobs, J., 1977, Rare earth and trace element geochemistry of metabasalts from the Point Sal ophiolite, California: Earth and Planetary Science Letters, v. 37, p. 203–215, doi: 10.1016/0012821X(77)90165-0. Menzies, M., Seyfried, W., and Blanchard, D., 1979, Experimental evidence of rare earth element immobility in greenstones: Nature, v. 282, p. 398–399, doi: 10.1038/282398a0. Morris, J.D., and Hart, S.R., 1983, Isotopic and incompatible element constraints on the genesis of island arc volcanics from Cold Bay and Amak Island, Aleutians, and implications for mantle structure: Geochimica et Cosmochimica Acta, v. 47, p. 2015–2030, doi: 10.1016/0016-7037(83)90217-X. Mullen, E.D., 1983, MnO/TiO2/P2O5: A minor element discriminant for basaltic rocks of oceanic environments and its implications for petrogenesis: Earth and Planetary Science Letters, v. 62, p. 53–62. Nagy, E.A., Samson, S.D., and D’Lemos, R.S., 2002, U-Pb geochronological constraints on the timing of Brioverian sedimentation and regional deformation in the St. Brieuc region of the Neoproterozoic Cadomian orogen: Precambrian Research, v. 116, p. 1–17, doi: 10.1016/S03019268(01)00235-2. Nance, R.D., and Murphy, J.B., 1994, Contrasting basement isotopic signatures and the palinspastic restoration of peripheral orogens: Example from the Neoproterozoic Avalonian-Cadomian belt: Geology, v. 22, p. 617–620, doi: 10.1130/0091-7613(1994)022<0617:CBISAT>2.3.CO;2. Ordoñez Casado, B., 1998, Geochronological studies of the Pre-Mesozoic basement of the Iberian Massif: The Ossa Morena zone and the allochtonous complexes within the Central Iberian zone [Ph.D. thesis]: Zurich, Eidgenossische Technische Hochschule Zürich, 235 p. Pacltová, B., 2000, Thalloid algae with reproductive organs and other fossils in limestones and solid bitumens: Neoproterozoic, Czech Republic: Veˇstnik Cˇeského geologického ustavu, v. 75, p. 339–350. Pasava, J., 2000, Normal versus metal-rich black shales in the Barrandian Neoproterozoic of the Teplá-Barrandian unit: A summary with new data: Veˇstnik Cˇeského geologického ustavu, v. 75, p. 229–239. Pearce, J.A., and Parkinson, I.J., 1993, Trace element models for mantle melting: Applications to volcanic arc petrogenesis, in Prichard, H.M., Alabaster, T., Harris, N.B.W., and Neary, C.R., eds., Magmatic processes and plate tectonics: London, Geological Society of London Special Publication 76, p. 373–403. Pearce, J.A., and Peate, D.W., 1995, Tectonic implications of the composition of volcanic arc magmas: Annual Review of Earth and Planetary Science, v. 23, p. 251–285, doi: 10.1146/annurev.ea.23.050195.001343. Pearce, J.A., Ernewein, M., Bloomer, S.H., Parson, L.M., Murton, B.J., and Johnson, L.E., 1995, Geochemistry of Lau Basin volcanic rocks: Influence of ridge segmentation and arc proximity, in Smellie, J.L., ed., Volcanism associated with extension at consuming plate margins: London, Geological Society of London Special Publication 81, p. 53–75. Pelc, Z., and Waldhausrová, J., 1994, Geochemical characteristics of volcanites of the Barrandian Upper Proterozoic in the Chudenice-Nepomuk region, SW Bohemia: Sborník Geologických Veˇd, Geologie v. 46, p. 5–21. Pin, C., and Paquette, J.L., 1997, A mantle-derived bimodal suite in the Hercynian belt: Nd isotope and trace element evidence for a subduction-related rift origin of the late Devonian Brévenne metavolcanics, Massif Central (France): Contributions to Mineralogy and Petrology, v. 129, p. 222–238, doi: 10.1007/s004100050334. Pin, C., and Santos Zalduegui, J.F., 1997, Sequential separation of light rareearth elements, thorium and uranium by miniaturized extraction chromatography: Application to isotopic analyses of silicate rocks: Analytica Chimica Acta, v. 339, p. 79–89, doi: 10.1016/S0003-2670(96)00499-0. Pin, C., Liñán, L., Pascual, E., Donaire, T., and Valenzuela, A., 2002, Late Proterozoic crustal growth in the European Variscides: Nd isotope and geochemical evidence from the Sierra de Cordoba andesites (Ossa Morena zone, Southern Spain): Tectonophysics, v. 352, p. 133–151. Pouba, Z., Krˇibek, B., and Pudilová, M., 2000, Stromatolite-like cherts in the Barrandian Upper Proterozoic: A review: Veˇstnik Cˇeského geologického ustavu, v. 75, p. 285–296.
Sm-Nd isotope and trace element study of Late Proterozoic metabasalts Pouclet, A., Lee, J.S., Vidal, P., Cousens, B., and Bellon, H., 1995, Cretaceous to Cenozoic volcanism in South Korea and in the Sea of Japan: Magmatic constraints on the opening of the back-arc basin, in Smellie, J.D., ed., Volcanism associated with extension at consuming plate margins: London, Geological Society of London Special Publication 81, p. 169–191. Price, R.C., Johnson, L.E., and Crawford, A.J., 1990, Basalts of the North Fiji Basin: The generation of back-arc basin magmas by mixing of depleted and enriched mantle sources: Contributions to Mineralogy and Petrology, v. 105, p. 106–121, doi: 10.1007/BF00320970. Prouteau, G., Scaillet, B., Pichavant, M., and Maury, R., 2001, Evidence for mantle metasomatism by hydrous silicic melts derived from subducted oceanic crust: Nature, v. 410, p. 197–200, doi: 10.1038/35065583. Quesada, C., 1990, Precambrian successions in SW Iberia: Their relationship to “Cadomian” orogenic events, in D’Lemos, R.S., Strachan, R.A., and Topley, C.G., eds., The Cadomian orogen: London, Geological Society of London Special Publication 51, p. 353–362. Quesada, C., Apalategui, O., Eguiluz, L., Liñan, E., and Palacios, T., 1990, Stratigraphy. Precambrian, in Dallmeyer, R.D., and Martinez Garcia, E., eds., Pre-Mesozoic geology of Iberia: Berlin, Springer-Verlag, p. 252–258. Röhlich, P., 2000, Some stratigraphic problems of the Barrandian Neoproterozoic: Veˇstnik Cˇeského geologického ustavu, v. 75, p. 201–204. Samson, S.D., D’Lemos, R.S., Blichert-Toft, J., and Vervoort, J., 2003, U-Pb geochronology and Hf-Nd isotope compositions of the oldest Neoproterozoic crust within the Cadomian orogen: New evidence for a unique juvenile terrane: Earth and Planetary Science Letters, v. 208, p. 165–180, doi: 10.1016/S0012-821X(03)00045-1. Saunders, A.D., and Tarney, J., 1984, Geochemical characteristics of basaltic volcanism within back-arc basins, in Kokelaar, B.P., and Howell, M.F., eds., Marginal basin geology: Oxford, Blackwell Scientific Publications for the Geological Society of London, p. 59–76. Saunders, A.D., Tarney, J., Marsh, N.G., and Wood, D.A., 1980, Ophiolites as an ocean crust or marginal basin crust: A geochemical approach, in Panayiotou, A., ed., Ophiolites. Proceedings of the International Ophiolite Symposium, Cyprus, 1979: Cyprus, Geological Survey Department of Cyprus, p. 193–204. Schäfer, H.J., Gebauer, D., Nägler, T.F., and Eguiluz, L., 1993, Conventional and ion-microprobe U-Pb dating of detrital zircons of the Tentudia Group (Serie Negra, SW Spain): Implications for zircon systematics, stratigraphy, tectonics and the Precambrian/Cambrian boundary: Contributions to Mineralogy and Petrology, v. 113, p. 289–299, doi: 10.1007/BF00286922. Schellart, W.P., 2004, Kinematics of subduction and subduction-induced flow in the upper mantle: Journal of Geophysical Research, v. 109, p. B07401, doi: 10.1029/2004JB002970. Shervais, J., 1982, Ti-V pots and the petrogenesis of modern ophiolitic lavas: Earth and Planetary Science Letters, v. 59, p. 101–118. Sinton, J.M., and Fryer, P., 1987, Mariana Trough basalts from 18°N: Implications for the origin of back arc basin basalts: Journal of Geophysical Research, v. 92, p. 12,782–12,802. Skocˇek, V., and Pouba, Z., 2000, Neoproterozoic sedimentary carbonates and their silicified equivalents: Barrandian, Czech Republic: Veˇstnik Cˇeského geologického ustavu, v. 75, p. 241–260. Staudigel, H., and Hart, S.R., 1983, Alteration of basaltic glass: Mechanisms and significance for the oceanic crust-seawater budget: Geochimica et Cosmochimica Acta, v. 47, p. 337–350, doi: 10.1016/0016-7037(83)90257-0. Sun, S.S., 1980, Lead isotopic study of young volcanic rocks from mid-ocean ridges, ocean islands and island arcs: Philosophical Transactions of the Royal Society of London, v. A297, p. 409–445. Sun, S.S., and McDonough, W.F., 1989, Chemical and isotopic systematics of oceanic basalts: Implications for mantle composition and processes, in
247
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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Figure 1. Geological map of the Prague synform showing the main structural features; locations of the major mapped faults; and Cretaceous, Tertiary, and Silurian outcrops. The studied outcrop is indicated by the circled number 1: Svatý Jan pod Skalou.
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Figure 2. Detailed geological map with a zoom on the two sampled dikes (circled numbers 1 and 2).
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Figure 3. (A) Block diagram representing both dike1 and dike2. 1—alkaline basaltic dikes (dike1 may be composed of two successive intrusions); 2—thin bedded limestone alternating with calcareous shale; 3—largely hyaloclastic alkaline-basaltic tuff with common volcanic bombs; 4—sampled sections across dike1 and dike2, respectively; 5—feldspar phenocrysts aligned along the contact rim of dike2. (B) Magnetic lineations (K1) from the borders and the center of the dikes showing the recorded magnetic fabric. All images are equal-area mappings. Triangles—mean vector; squares—eigenvectors of the cylindrical best fit; dashed lines—dike trace after unfolding; n—number of specimens used for the statistics. Contour intervals: 1 at a significance level of ±1σ.
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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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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.
d1/C1
d1/C2
d2/C2 d2/C2 d1/C2
d1/C1 36V
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
B
A
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C
α
N
NW K1
1 W
B
Rheic
K1
E
α
K1
K1
2
1a SE K1 K1
G Dike1
"Opening Mode"
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.
Jelen´ska, M., Aïfa, T., Ka˛działko-Hofmokl, M., and Zelazniewicz, A., 2001, Paleomagnetic and rock magnetic study of Devonian olistoliths set in Early Carboniferous flysch, West Sudetes, Poland: Geophysical Journal International, v. 145, p. 821–834, doi: 10.1046/j.0956-540x.2001.01448.x. Jelínek, V., 1978, Statistical processing of anisotropy of magnetic susceptibility measured on groups of specimens: Studia Geophysica & Geodaetica, v. 22, p. 50–62, doi: 10.1007/BF01613632. Jelínek, V., and Pokorný, J., 1997, Some new concepts in technology of transformer bridges for measuring susceptibility anisotropy of rocks: Physics and Chemistry of the Earth, v. 22, p. 179–181, doi: 10.1016/S00791946(97)00099-2. Kachlík, V., and Patocˇka, F., 1998, Cambian/Ordovician intracontinental rifting and Devonian closure of the rifting generated basins in the Bohemian Massif realms: Acta Universitatis Carolinae-Geologica, v. 42, no. 3/4, p. 433–441. Kamb, W.B., 1959, Ice petrofabric observations from Blue Glacier, Washington, in relation to theory and experiments: Journal of Geophysical Research, v. 64, p. 1891–1909. Keppie, J.B., Dostal, J., Dallmeyer, R.D., and Doig, R., 2000, Superposed Neoproterozoic and Silurian magmatic arcs in central Cape Breton Island, Canada: Geochemical and geochronological constraints: Geological Magazine, v. 137, p. 137–153, doi: 10.1017/S0016756800003769. Kirschvink, J.L., 1980, The least squares line and plane and the analysis of paleomagnetic data: Geophysical Journal of the Royal Astronomical Society, v. 62, p. 699–718. Krˇíž, J., 1998, Silurian, in Chlupácˇ, I., Havlícˇek, V., Krˇíž, J., Kukal, Z., and Štorch, P., 1998, Palaeozoic of the Barrandian (Cambrian to Devonian): Prague, Czech Geological Survey, p. 79–101. Kröner, U., and Hahn, T., 2003. Sedimentation, Deformation und Metamorphose im Saxothuringikum während der variszsichen Orogenese: Die komplexe Entwicklung von Nord-Gondwana während kontinentaler Subduktion und schiefer Kollision, in Linneman, U., ed., Das Saxothuringikum: Dresden, Staatliche Naturhistorische Sammlungen Dresden, p. 133–146. Krs, M., Pruner, P., and Man, O., 2001, Tectonic and paleogeographic interpretation of the paleomagnetism of Variscan and pre-Variscan formations of the Bohemian Massif, with special reference to the Barrandian terrane: Tectonophysics, v. 332, no. 1–2, p. 93–114, doi: 10.1016/S00401951(00)00251-1. Kukal, Z., and Jäger, O., 1988, Siliciclastic signal of the Variscan orogenesis: The Devonian Srbsko Formation of central Bohemia: Vestnik Ustedniho ustavu geologickeho, v. 63, p. 65–81. Lefort, J.P., Aïfa, T., and Hervé, F., 2006, Structural and AMS study of a Miocene dike swarm located above the Patagonian subduction, in Hanski, E., Mertanen, S., Rämö, T., and Vuollo, J., eds., Dyke swarms—Time markers of crustal evolution: London, Taylor & Francis, p. 225–241. Lewandowski, M., 1997, Palaeomagnetism of early Palaeozoic terranes between Baltica and Avalonia. EUROPROBE meeting. TESZ project, Potsdam. Extended abstract: Terra Nostra, v. 11, p. 69–72. Lewandowski, M., 1998. Assembly of Pangea: Combined palaeomagnetic and palaeoclimatic approach, in Ginter, M., Wilson, M.V.H., eds., CircumArctic Palaeozoic faunas and facies, IGCP406: Warsaw, Ichthyolith Issues Special Publication 4, p. 29–32. Lewandowski, M., 1999, A palaeomagnetic study of fracture fills in the Holy Cross Mountains of central Poland and its application in dating tectonic processes: Geophysical Journal International, v. 137, p. 783–792. Linnemann, U., McNaughton, N.J., Romer, R.L., Gehmlich, M., Drost, K., and Tonk, C., 2004, West African provenance for Saxo-Thuringia (Bohemia Massif): Did Armorica ever leave pre-Pangean Gondwana? U/PbSHRIMP zircon evidence and the Nd-isotopic record, Germany: International Journal of Earth Sciences, v. 93, p. 683–705, doi: 10.1007/s00531004-0413-8. Man, O., 2003, Analysis of multi-component natural remanent magnetization based on the thermal demagnetization spectrum: Studia Geophysica & Geodaetica, v. 47, p. 359–370, doi: 10.1023/A:1023727809339. Marheine, D., Kachlík, V., Patocˇ ka, F., Maluski, H., and Zelazniewicz, A., 2000, Variscan polyphase tectonothermal record in the West Sudetes (Bohemian Massif)—Deduced from Ar-Ar ages, in Galicia 2000 basement tectonics 15: Variscan–Appalachian dynamics: The building of the Upper Paleozoic basement, 15th International Conference on Basement Tectonics, Galicia, La Coruna, Spain, 4–8 July: International Basement Tectonics Association, p. 254–257.
Structural evolution of the Prague synform (Czech Republic) during Silurian times Matte, P., 1991, Accretionary history and crustal evolution of the Variscan Belt in western Europe: Tectonophysics, v. 196, p. 309–339. Matte, P., 2001, The Variscan collage and orogeny (480–290 Ma) and the tectonic definition of the Armorica microplate: A review: Terra Nova, v. 13, p. 122–128, doi: 10.1046/j.1365-3121.2001.00327.x. Matte, P., Maluski, H., Rajlich, P., and Franke, W., 1990, Terrane boundaries in the Bohemian Massif: Result of a large-scale Variscan shearing: Tectonophysics, v. 177, p. 151–170, doi: 10.1016/0040-1951(90)90279-H. McKerrow, W.S., Mac Niocaill, C., and Dewey, J.F., 2000, The Caledonian orogeny redefined: Journal of the Geological Society of London, v. 157, no. 6, p. 1149–1154. Melichar, R., 2004, Tectonics of the Prague synform: A hundred years of scientific discussion: Krystalinikum, v. 30, p. 167–187. Moreira, M., Geoffroy, L., and Pozzi, J.P., 1999, Ecoulement magmatique dans les dykes du point chaud des Açores: étude préliminaire par anisotropie de susceptibilité magnétique (ASM) dans l’île de San Jorge: Comptes Rendus de l’Académie des Sciences, v. 329, p. 15–22. Murphy, J.B., and Nance, R.D., 2003, Do supercontinents introvert or extrovert? Sm-Nd isotope evidence: Geology, v. 31, p. 873–876, doi: 10.1130/ G19668.1. Parma, J., and Zapletal, K., 1991, CS-1 apparatus for measuring the temperature dependence of low-field susceptibility of minerals and rocks (in cooperation with the KLY-2 Kappabridge). Leaflet: Brno, Czech Republic, Geofyzika Brno, 24 p. Patocˇka, F., Vlašímsky, P., and Blechová, K., 1993, Geochemistry of Early Paleozoic volcanics of the Barrandian basin (Bohemian Massif, Czech Republic): Implications for Paleotectonic reconstructions: Wien, Jahrbuch der Geologischen Bundesanstalt, v. 136, p. 871–894. Patocˇka, F., Pruner, P., and Štorch, P., 2003, Palaeomagnetism and geochemistry of Early Paleozoic rocks of the Barrandian (Teplá-Barrandian unit, Bohemian Massif): Paleotectonic implications: Physics and Chemistry of the Earth, v. 28, p. 735–749. Philpotts, A.R., and Asher, P.M., 1994, Magmatic flow-direction indicators in a giant diabase feeder dyke, Connecticut: Geology, v. 22, p. 363–366, doi: 10.1130/0091-7613(1994)022<0363:MFDIIA>2.3.CO;2. Prˇíhoda, K., Krs, M., Pešina, B., and Bláha, J., 1989, MAVACS—A new system of creating a nonmagnetic environment for palaeomagnetic studies: Cuadernos de Geologica Ibérica, v. 12, p. 223–250. Puranen, R., Pekkarinen, L.J., and Pesonen, L.J., 1992, Interpretation of magnetic fabrics in the Early Proterozoic diabase dykes of Keuruu, central
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Finland: Physics of the Earth and Planetary Interiors, v. 72, p. 68–82, doi: 10.1016/0031-9201(92)90050-6. Robardet, M., 2003, The Armorica “microplate”: Fact or fiction? Critical review of the concept and contradictory palaeogeographical data: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 195, p. 125–148, doi: 10.1016/ S0031-0182(03)00305-5. Schätz, M., Reischmann, T., Tait, J., Bachtadse, V., Bahlburg, H., and Martin, U., 2002, The Early Palaeozoic break-up of northern Gondwana, new palaeomagnetic and geochronological data from the Saxothuringian basin, Germany: International Journal of Earth Sciences, v. 91, p. 838–849, doi: 10.1007/s00531-002-0261-3. Schulz, B., Bombach, K., Pawlig, S., and Brätz, H., 2004, Neoproterozoic to Early Palaeozoic magmatic evolution in the Gondwana-derived Austroalpine basement to the south of the Tauern Window (Eastern Alps): International Journal of Earth Sciences, v. 93, no. 5, p. 824–843, doi: 10.1007/ s00531-004-0421-8. Smith, J.V., Miyake, Y., and Yamauchi, S., 1993, Flow direction and groundmass shear zones in dykes, Shimane Peninsula, Japan: Geological Magazine, v. 130, no. 1, p. 117–120. Stesky, R.M., and Pearce, G.W., 1995, Spheristat 2.2 for Windows: Brockville, Ontario, Canada, Pangaea Scientific, 199 p. Štorch, P., 1987, Silur, in Havlícˇek, V., ed., 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. Štorch, P., 1998, Volcanism, in Chlupácˇ, I., Havlícˇek, V., Krˇíž, J., Kukal, Z., and Štorch, P., eds., Palaeozoic of the Barrandian (Cambrian to Devonian): Prague, Czech Geological Survey, p. 149–163. Torsvik, T.H., and Cocks, L.R.M., 2004, Earth geography from 400 to 250 Ma: A palaeomagnetic, faunal and facies review: Journal of the Geological Society of London, v. 161, p. 555–572. Von Raumer, J.F., Stampfli, G.M., and Bussy, F., 2003, Gondwana-derived microcontinents: The constituents of the Variscan and Alpine collisional orogens: Tectonophysics, v. 365, p. 7–22, doi: 10.1016/S0040-1951(03)00015-5. Žák, J., Schulmann, K., and Hrouda, F., 2005, Multiple magmatic fabrics in the Sázava pluton (Bohemian Massif, Czech Republic): A result of superposition of wrench-dominated regional transpression on final emplacement: Journal of Structural Geology, v. 27, p. 805–822, doi: 10.1016/ j.jsg.2005.01.012. MANUSCRIPT ACCEPTED BY THE SOCIETY 3 OCTOBER 2006
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 <5% (based on 1σ standard deviation) for the rare earth elements (REE), Hf, Nb, and Ta, except for Sm, Lu, Hf, and Th (5–13%). Reproducibility, accuracy, and precision of the applied method were controlled by repeated analysis of NIST SRM 614 (data by Horn et al., 1997) and whole-rock geostandards (BE-N Basalt, MAG-1 Marine Mud; Govindaraju, 1994). Based on analysis of NIST SRM 614, the precision for REE, Nb, Ta, and Th is <5% (based on 1σ), except for Gd and Er (5–7.5%). The mineral chemistry (major elements) has been analyzed by EMP (900 points, CAMECA SX 50, SX 51 at 15 kV, 10 nA, counting time 20 s, with CAMECA routine ZAF and PAP corrections; for representative data see Table 1) in eleven micaschist and garnet gneiss samples and in five amphibole-bearing metabasites. Amphibole was characterized by profiles of five to fifteen points. The cores and rims of chlorites, epidotes, and plagioclases next to the amphibole porphyroblasts were analyzed and further profiles with three to five points were taken. Cations in amphiboles were normalized to 23 oxygens and sum(T1 + T2 + M1 + M2 + M3) = sumMg = 13 (13eCNK), as detailed in Triboulet (1992). Fe3+ has been estimated as maximum. In micaschists and paragneisses, the composition of zoned garnet, staurolite, cordierite, biotite, muscovite, and plagioclase was determined from cores to rims. The core of mica inclusions in garnet was also analyzed. The chemical evolution of garnets in micaschists was studied in several samples by profiles with a point spacing between 0.05 and 0.2 mm. However, single profiles may not have passed the entire core region; furthermore, porphyroblasts display zonation gaps, and some show only a part of the complete chemical evolution of garnet. These complications are not necessarily evident from the single zonation profiles. Therefore, garnet compositions were investigated using spessartine-grossularpyrope coordinates. As the Mn component is often controlled by Rayleigh fractionation during crystallization (Hollister, 1966), it allows recognition of the relative temporal chemical growth evolution when garnet is the main Mn fractionating phase. Trace element analyses of garnet in monazite-bearing samples were carried out by LA-ICP-MS (Longerich et al., 1993) in those thin sections that were used for the EMP analyses. Garnet profiles included five to ten single-shot craters with diameters of 35 μm. The laser was adjusted to a scan speed of 5 μm/s, an energy of 0.23–0.28 mJ (40 J/cm2), and a repetition rate of 5 Hz. After 25 s
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of background detection, an analysis of 50 s with time-resolved signals and integration times of 10 ms for 29Si and 49Ti; 20 ms for 45 Sc, 51V, 53Cr, 89Y; and 30 ms for 7Li, 60Ni, 69Ga, and REE with masses 139–175 was compared to the glass reference material standards NIST SRM 612 and SRM 614 (Jackson et al., 1992; Pearce et al., 1997). SiO2 of garnet rims known from EMP analysis was used as the internal standard; an in-house standard garnet K23 allowed verification of the data. After signal quantification by the GLITTER 3.0, the 1σ errors, based on counting statistics from signal and background, range ~29% (Gd, 3.2 ± 1.1 ppm) and 14% (Yb, 123 ± 18 ppm), depending on the absolute concentrations. Averaged errors are given in Table 1. Ti in the zoning profiles has been considered, as no high Ti resulting from unintentional analysis of tiny Ti-phase inclusions in garnet occurred. The in situ “chemical” Th-U-Pb dating of monazite by EMP analysis (Montel et al., 1994, 1996; Suzuki et al., 1994) is based on the observation that concentration of common lead in monazite (light REE [LREE], Th)PO4 is negligible when compared to radiogenic lead resulting from decay of Th and U (Cocherie et al., 1998). As Th concentrations are mostly high (3–14 wt%), a sufficient amount of radiogenic lead to be analyzed by EMP can accumulate in Paleozoic monazite. All monazite analyses of this study were carried out on a JEOL JX 8600 at Salzburg University, using conditions of 15 kV, 250 nA, and a beam diameter of ~5 μm (Finger and Helmy, 1998). Mα1 lines were chosen for Th, U, and Pb; Lα1 for La, Y, and Ce; Lβ1 for Pr and Nd; and Kα1 for P, Si, and Ca. The counting times for Pb, Th, and U were 240–360, 30, and 50 s, respectively, per analysis on peak and 2 × 120, 2 × 15, and 2 × 25 s, respectively, on background. This procedure results in statistical errors (1σ) of typically 0.012, 0.05, and 0.015 wt% for Pb, Th, and U, respectively (Finger and Helmy, 1998). All other elements were determined with 10 s (2 × 5 s) counting times. The small Y interference on the Pb Mα line was corrected by linear extrapolation after measuring a Pbfree yttrium standard (Montel et al., 1996). A small Th interference on U Mα was also empirically corrected. For each single analysis, a chemical age was calculated using the equations of Montel et al. (1996) plus a respective error resulting from counting statistics, which was mostly between ±20 and ±40 m.y. (1σ). Weighted average ages were calculated after Ludwig (2001). To control the quality of analyses, monazite with a concordant U-Pb age of 341 ± 2 Ma analyzed by thermal ionization mass spectrometry (TIMS) (Friedl, 1997) has been measured together with the specimen. Weighted average model ages of 346 ± 13, 341 ± 9, 340 ± 8, and 344 ± 11 Ma were obtained from the standard during four different sessions and are in agreement with the U-Pb TIMS age. METAMORPHIC EVOLUTION Bulk Rock Compositions of Metasediments For evaluation of bulk rock compositions, monazite-bearing metasediments can be subdivided into micaschists with garnet and staurolite, paragneisses without garnet, and aluminous paragneisses
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with garnet. Metagranitoids, such as the Gneiss de Brest, the Pte des Renards granite gneiss, and the ortho-augengneiss of Lochrist, were analyzed for comparison. In the La-Th-Sc and Th-Sc-Zr/10 discrimination diagrams for the provenance character and geotetonic setting (Bhatia and Crook, 1986), all former clastic sediments (micaschists and paragneisses) as well as the Léon metagranitoids uniformly plot in the field of a continental magmatic arc (Fig. 2A). In the Ni-TiO2 and SiO2-K2O/Na2O discrimination diagrams of Floyd et al. (1989) and Roser and Korsch (1986), the provenance characteristics are confirmed by indication of a source region dominated by felsic rocks and by the placement of the sequence in an active continental margin setting (Fig. 2B). It should be noted that the analyzed rocks have a Ca-poor composition, which is typical for monazite-bearing metasediments. Nethertheless, the data are regarded as being representative for the recognition of the sedimentary source region, because the lithological units are rather monotonous. Monazite-bearing rock types appear to be abundant within the suite. Similar geochemical
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Figure 2. Bulk rock compositions of monazite-bearing metasediments from the Léon domain (A) in the diagrams La-Sc-Th and Th-Sc-Zr/10 (Bhatia and Crook, 1986) and (B) log (K2O/Na2O)-SiO2 (Roser and Korsch, 1986). Discriminative compositional fields: ACM—active continental margin; CMA—continental magmatic arc; IA—island arc; OIA—oceanic island arc; PCM—passive continental margin. Compositions of orthogneisses of Brest and Plounevez-Lochrist are shown for comparison.
characteristics and geotectonic setting have been described from other Neoproterozoic metasedimentary sequences in the SaxoThuringian zone (Linnemann and Romer, 2002). Mineral Chemistry and Geothermobarometry of Metabasites Lenses of eclogites and their retrogressed equivalents occur around the contact of paragneisses and the ortho-augengneiss of Plounevez-Lochrist in the lower parts of the Lesneven Gneiss unit. In the samples from location Kao (in the village of Plounevez-Lochrist), with the best preserved eclogites, garnet of a relict eclogitic peak metamorphic assemblage displays a prograde zonation, with increasing pyrope contents (20–45 wt%) and decreasing grossular (30–18 wt%), almandine (47–35 wt%), and spessartine (3.3–0.5 wt%) contents from core to rim. As was similarly observed by Paquette et al. (1987) and Godard and Mabit (1998), no clinopyroxene was preserved in the matrix, and symplectites with plagioclase and amphibole prevail. The garnet core composition corresponds to the observations by Paquette et al. (1987), who calculated 700 °C at 13–14 kbar using garnet and enclosed omphacite with Jd 27% by garnet-clinopyroxene equilibria (Ellis and Green, 1979; Krogh, 1988) and related geobarometers (Holland, 1983). This P-T estimate represents the minimum temperatures as well as corresponding minimum pressures that occurred during the crystallization of garnet cores (see Fig. 5A). Although the increase of Mg in garnet can be attributed to growth during increasing temperature, decrease of Ca toward the garnet rim signalizes decreasing pressure during crystallization of the garnet-bearing assemblage. Godard and Mabit (1998) found further petrological evidence for such a temperaturedominated evolution subsequent to the peak pressure conditions. Ca-amphibole, which replaced garnet and clinopyroxene in overprinted eclogites, is also abundant in the numerous meterscale metabasite lenses and in the large Lannilis metabasite complex (Fig. 1B). Ca-amphibole is best described by its IVAl-VIAl and Ti-IVAl composition, but the other chemical parameters, Si (decreases with increase of Altot), Mg, NaM4, and (Na + K)A show the corresponding variations (Triboulet and Audren, 1988). P-T conditions of the amphibole-bearing assemblages were estimated by a least squares approximation of amphibole compositional dependence on temperature and pressure for the assemblage Caamphibole + plagioclase (An > 10) + quartz ± zoisite ± chlorite ± calcite, determined by Plyusnina (1982), as used by Gerya et al. (1997) and empirically modified for Fe3+ in amphibole by Zenk and Schulz (2004): T[K] = 4701/[1.825 – 1.987 ln(8/15.5 – SiAm/sumKat + 0.07531)] P[kbar] = ([−425 – 1719 (XAlM/XAlM + [Fe3+ /2.763]) + 2.75T Am + 1.987T ln(XAlM)] + 1)/1000, with T in kelvins; XAlM = (SiAm + AlAm – 8)/2.763; and SiAm, AlAm, Fe3+ , sumKat are Si, Altot, Fe3+, and sum of cations in amphibole. Am
Cadomian and Variscan metamorphic events in the Léon domain In the retrogressive post-peak assemblage of the kyanitebearing eclogite (sample Kao), unzoned magnesio-hornblende with IVAl 1.0/VIAl 0.3 shows no preferential orientations. It coexists with andesine, epidote, quartz, and sphene and crystallized at 530 °C and 4 kbar (Fig. 5A). Preferentially oriented tschermakite in an amphibolite from the Lannilis metabasite complex (sample Lan) in the upper part of the Lesneven Gneiss unit crystallized with andesine, epidote, quartz, and sphene and is unzoned, with IVAl 1.7/VIAl 0.4 and elevated Ti of 0.13 compared to the other samples. These amphiboles indicate metamorphic conditions at 620 °C and 6 kbar, when the thermobarometer reported above is used. An amphibolite with a similar mineral assemblage (sample Kerhorn) from the western part of the Lesneven Gneiss unit displays zonations from magnesio-hornblende (IVAl 1.2/VIAl 0.5) in the cores to actinolite rims (IVAl 0.7/VIAl 0.2). The continuous zonation corresponds to a retrograde P-T path from 540 °C and 5 kbar to 400 °C and 2 kbar (Fig. 5A). Amphibolite samples from the Conquet-Penze Micaschist unit come from Pors Liogan (sample PLi3 in Fig. 5) and the harbor of Le Conquet (PortCo). Long axes of preferentially oriented green amphiboles define a southwest-plunging mineral lineation, associated with a weak foliation that is parallel to the structural elements in surrounding micaschists. The unzoned tschermakitic hornblendes and tschermakites (Leake et al., 1997) with IVAl 1.8/ VI Al 1.0 in sample PLi3 and tschermakitic amphiboles with IVAl 1.5/VIAl 0.45 in sample PortCo coexist with andesine, epidote, quartz, and spene. Garnet occurs in sample PLi3. The amphiboles recorded 610 °C at 8 kbar in sample PLi3 and 590 °C at 6 kbar in sample PortCo and are interpreted to represent the conditions of the thermal maximum in this unit (Fig. 5B). Mineral Chemistry and Geothermobarometry of Garnet Micaschists and Paragneisses Lesneven Gneiss Unit Garnet is rare in the paragneisses of the Lesneven Gneiss unit. The distinct 1.5-km-long northeast-striking horizon with aluminous paragneisses in the vicinity of the eclogite lenses to the south of Plounevez-Lochrist (Chauris et al., 1998) seems to be the only garnet-bearing paragneiss location in the lower part of this unit. In samples from La Garenne (Lage), Traonjulien (Trao), and Keranton (Kerz), a foliated matrix by biotite (muscovite only rarely appears as a late phase), plagioclase, K-feldspar, quartz, and both kyanite and fibrolitic sillimanite surround garnet (up to 4 mm in diameter). Cordierite is abundant in the matrix of samples Trao (Fig. 3A) and Kerz. As kyanite is enclosed, cordierite is interpreted to have crystallized after kyanite. Andalusite appears as a late phase in sample Trao. Cordierite in sample Kerz has higher Na2O (0.55–0.70 wt%) than in sample Trao (0.40–0.50 wt%), whereas the XMg of cordierite is similar in both samples (XMg 0.73–0.78), with slightly higher XMg in the cores. The poikiloblastic garnet encloses biotite, plagioclase, quartz, and kyanite. Chemical analysis along profiles (Fig. 4) indicates that the porphyroblasts have
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large homogeneous cores with 25% pyrope, 8% grossular, and 5% spessartine contents. In the narrow outermost rim, Mg significantly decreases while Fe increases at almost constant Ca and Mn (Fig. 4A, H, and I), indicating a temperature decrease. The XMg of biotite enclosed in garnet and in the matrix are similar and depend on the XMg in garnet cores. Plagioclase is andesine and zoned with decreasing anorthite contents toward the rim. The P-T conditions for crystallization of garnet cores have been calculated using the garnet-biotite thermometer of Bhattacharya et al. (1992) in combination with the garnet-sillimanite-plagioclase (GASP) barometer of Holland and Powell (1990), involving an internally consistent thermodynamic data set with the ideal activity models for garnet (Ganguly and Saxena, 1984; Ganguly et al., 1996) and plagioclase (Powell and Holland, 1993). Furthermore, the garnet-cordierite thermobarometer using the thermodynamic data set by Perchuk (1991) was also applied. Tentative calculations by other calibrations of the thermo- and barometers yielded no substantially different results. Maximum temperatures calculated from the Mg-rich garnet cores and enclosed biotite or matrix biotite with a similar XMg range between 740 and 780 °C. Corresponding pressures of 8–10 kbar were calculated by plagioclase enclosed in garnet or from the matrix (Fig. 5A). Temperatures and pressures from the garnet-cordierite equilibria are ~800 °C at 7 kbar. This difference is in the range of the general error of the thermobarometric estimates and can be attributed to the different thermodynamic data sets involved. According to the P-T data, the lack of muscovite in the assemblage with garnet can be explained by its prograde decomposition to K-feldspar and sillimanite. Godard and Mabit (1998) attributed the appearance of sapphirine after kyanite in quartz-free eclogites nearby to a substantial increase of temperature during the decompression. This increase is further confirmed by the thermobarometric data from the aluminous paragneisses. In the upper part of the Lesneven Gneiss unit, which is exposed along the coast to the west (Fig. 1B), one sample of garnet gneiss has been gained from a 5-m-thick horizon at Porz Pabu (sample Pabu) in the Kermorvan peninsula, corresponding to the location of sample LC12 in Jones (1994). Layers with relict garnet in a foliated matrix with biotite, plagioclase, quartz, staurolite, and kyanite are embedded in domains with biotite, muscovite, plagioclase, quartz, and fibrolitic sillimanite, the latter occuring with biotite in the S2 foliation. Apparently garnet has been replaced by aggregates of andalusite. Observations and interpretations of phase relations are given in Jones (1993). The garnet porphyroblasts are poorly zoned, with homogeneous cores of pyrope 13.3% and grossular 2.5% (Fig. 4B and G). Increase of spessartine at decreasing pyrope in the outer rim is attributed to retrograde exchange reaction. Temperatures of 635 °C were calculated from Mg-rich garnet zones and enclosed biotite or matrix biotite. They are identical to the results of Jones (1994) and represent the thermal peak. Jones (1994) argued that GASP barometry with low-An plagioclase (An 9) may overestimate pressure, because of poorly understood activity-composition relationships (Ashworth and Evirgen, 1985). However, in regional studies
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Figure 3. Microstructures in metasediments. Arrows mark craters produced by LA-ICP-MS analysis of trace elements in garnet. (A) Aluminous paragneiss (Trao) from Lesneven Gneiss unit. (B) Micaschist (PLiN1, Porz Liogan) from the Conquet-Penze unit, with curved internal foliation S1i in large garnet and straight S2 overgrown by staurolite. (C) Micaschist (Penz, Penzer) from Conquet-Penze unit, with anastomozing foliation S2 by preferentially oriented second generation of mica. An early generation of mica is oriented at acute angles or perpendicular to S2. (D) Monazite partly enclosed by ilmenite in matrix of aluminous paragneiss (Trao). Bt—biotite; Cd—cordierite; Grt—garnet; Ilm—ilmenite; Ky—kyanite; Ms—muscovite; Mz—monazite; Pl—plagioclase; Qtz—quartz; Sta—staurolite.
involving comparison of samples with oligoclase and low-An plagioclase, Schulz et al. (1998) showed that the garnet-muscovite-biotite-plagioclase (GMBP) barometers can be alternatively applied to assemblages with low-An plagioclase. Accordingly, pressures of 6.4 kbar at 630 °C (6.8 kbar with GASP) have been calculated for sample Pabu (Fig. 5A). Conquet-Penze Micaschist Unit Garnet is abundant in the Conquet-Penze Micaschist unit. Samples for geothermobarometry and monazite dating come from the outcrops at Penzer (Penz), Plage de Pors Liogan (samples labeled with “PLi”: PLi412, PLiN1, PLi1–5-1), Pointe des Renards (RenS3, Bil), Plage de Portez (samples labeled with “Port”: Portez, PortMic). Micaschist horizons in the coastal outcrops to the south of the Le Conquet have garnet of 1.5–5 cm in diameter, and a complex structural evolution is documented in
S1–S5 internal and external foliations (Jones, 1994). In samples with garnet <1 cm and without aluminosilicates, an internal foliation, S1i, which is outlined by numerous aligned inclusions of quartz, ilmenite, and rarely plagioclase and mica, is enclosed in garnet (Fig. 3B). Apparently dependent on the orientation of the section, the trails of S1i can be almost straight, slightly curved in an S-shape, but also more complex, with millipede-like structures, as shown by Passchier and Trouw (1996). The S1i is discordant to and does not line up with the surrounding matrix foliation S2. The main foliation S2 (corresponding to S4 in Jones, 1994) curves around the lenticular domains with garnet porphyroblasts, and its pressure shadows with quartz. A composite character of S2 and its polyphase development is obvious from intrafolial planar domains and lenticular microlithons bearing mica oriented at acute angles or perpendicular to the strictly preferentially oriented phyllosilicates, which define the S2 planes (Fig. 3C). Fine-grained
Cadomian and Variscan metamorphic events in the Léon domain
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TABLE 1. MAJOR AND TRACE ELEMENT DATA OF ZONED GARNET IN MICASCHISTS AND PARAGNEISSES Lage Trao Pabu Penz PLiN1 PortMic Error Core Core Rim Core Rim Core Mid Rim Core Rim Core Mid Rim2 1σ Sample 5m-05 227-04 264-01 406-04 398-08 92-04 95-06 103-09 364-09 371-02 36-06 41-09 46-12 (%) (wt%) SiO2 38.16 37.04 36.86 36.90 36.87 36.87 36.94 36.90 36.87 36.51 36.89 37.31 37.34 Al2O3 21.66 20.81 20.68 20.32 20.42 20.89 20.84 21.31 20.34 20.73 20.72 20.54 20.82 FeO 30.16 32.28 33.52 34.07 33.50 32.73 33.52 36.82 32.25 35.86 32.80 34.66 36.38 MnO 1.90 2.26 5.05 3.80 4.99 5.05 3.78 0.20 3.63 0.16 5.45 3.95 1.01 MgO 6.57 5.19 2.79 3.27 2.64 1.98 2.19 3.20 2.37 3.60 2.29 2.82 3.45 CaO 2.12 1.97 1.84 0.86 0.90 2.98 2.93 2.22 2.84 2.14 1.97 1.62 1.28 Total 100.57 99.56 100.74 99.22 99.32 100.50 100.20 100.65 98.30 99.00 100.12 100.90 100.28 (pfu) Si 2.978 2.967 2.973 3.005 3.006 2.971 2.980 2.960 3.021 2.970 2.986 2.998 3.000 Al 1.993 1.965 1.966 1.949 1.962 1.985 1.982 2.015 1.964 1.989 1.977 1.945 1.972 Fe 1.970 2.163 2.261 2.320 2.285 2.208 2.258 2.464 2.209 2.441 2.221 2.329 2.445 Mn 0.126 0.154 0.345 2.626 0.345 0.345 0.259 0.014 0.251 0.010 0.374 0.269 0.069 Mg 0.766 0.621 0.335 0.397 0.321 0.239 0.264 0.383 0.289 0.435 0.277 0.338 0.413 Ca 0.178 0.169 0.159 0.075 0.079 0.258 0.254 0.191 0.249 0.185 0.171 0.140 0.110 (%) Alm 64.8 68.1 71.8 75.4 74.9 72.7 75.3 80.8 73.7 78.7 72.2 74.9 80.2 Prp 25.2 21.0 11.3 13.3 10.8 7.9 8.7 12.5 9.6 14.7 9.4 11.4 13.8 Sps 4.3 5.2 11.6 8.8 11.6 11.3 8.5 0.5 8.4 0.4 12.6 9.0 2.3 Grs 5.7 5.7 5.3 2.5 2.7 8.1 7.5 6.2 8.3 6.2 5.8 4.7 3.7 (ppm) Li 20.8 27.3 50.5 73.4 55.5 60.2 112 22.5 126 54.1 33.5 78.4 26.3 9.8 Sc 178 135 150 33 20 103 75 80 81 67 104 121 110 14.2 Ti 290 388 86 282 259 384 278 111 363 127 264 251 101 30.9 V 88 185 144 36 51 83 62 2.7 53 34 26 58 42 12.2 Cr 156 151 128 83 114 113 140 129 158 71 27 101 54 29.3 Ga 6.5 12.4 8.7 4.0 3.6 5.9 4.8 5.9 5.0 4.1 4.6 5.9 5.5 26.0 Y 304 213 256 162 18 482 1112 239 1222 370 169 1129 336 28.8 LREE 1.9 2.7 2.4 3.1 2.9 3.1 5.7 6.5 3.4 4.3 6.1 7.4 6.4 46.2 Gd 5.6 5.6 8.0 2.5 2.7 17.2 9.7 17.9 18.2 13.8 13.8 20.2 14.4 32.0 Tb 3.1 2.3 2.2 1.5 0.1 5.0 6.7 4.4 8.3 8.7 4.7 8.8 5.7 22.7 Dy 37.7 29.9 30.7 18.8 5.1 57.3 108.0 34.5 125.1 72.3 32.5 115.0 48.9 22.2 Ho 8.9 7.7 8.6 6.8 0.8 15.8 38.7 12.2 40.4 18.0 3.6 39.0 9.4 19.6 Er 32.2 26.3 29.9 23.5 1.5 50.8 163.1 37.8 166.0 44.4 9.2 141.0 26.9 22.8 Tm 4.7 4.5 3.5 4.4 0.3 9.0 32.5 4.2 28.9 4.0 1.1 25.0 3.6 20.7 Yb 27.2 30 24 36 0.6 69.8 237 35 232.0 21.6 8.9 195 23.3 23.9 Lu 2.9 4.3 2.1 5.5 0.5 9.4 30.0 4.3 37.4 5.6 0.6 28.6 2.4 22.2 Note: Analyzed by microprobe and laser ablation–inductively coupled plasma–mass spectrometer. Major elements are normalized to 12 oxygens. Alm—almandine; Grs—grossular; LREE—light rare earth elements; Prp—pyrope; Sps— spessartine. Core, Mid, and Rim refer to locations on the garnet samples.
staurolite is enclosed in garnet rims. Coarser grained staurolite in the matrix is preferentially oriented with its long axis parallel to the southwest-plunging mineral lineation. It encloses planar trails of quartz and ilmenite (S2i), which line up with the external S2 (Fig. 3B). The microstructures indicate a succession of assemblages (involving muscovite + quartz + plagioclase) with: 1. garnet + biotite ± chlorite, 2. garnet + biotite ± staurolite, and 3. staurolite ± biotite, which crystallized during the development of the S1-S1i and S2 foliations. In high strain and mylonite zones, the southeast-dipping main foliation S2 was overprinted by shear bands, which cut at acute angles. A uniform top-to-southwest–directed normal
sense of shear parallel to the mineral lineation and corresponding to the dextral strike-slip movement along the Léon shear zone described by Balé and Brun (1986) was observed. The metapelite garnets in the Conquet-Penze unit display prograde zonations (Tracy, 1982) with uniformly decreasing spessartine and increasing pyrope and almandine contents from core to rim. Retrogressive overprints in the outermost garnet rims, which show a decreasing pyrope content, are only poorly developed. In samples from Penzer and Porz Liogan to the south, grossular contents (Grs 6.5–9.2%) are slightly higher than in Plage de Portez (3.3–6.4%) to the north (Fig. 4C–F). This difference is matched by different An contents in plagioclase, with higher An contents to the south (An 11–24) than to the north (An 8–15), as has similarly
276
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A
B
Lage
Pabu
1 mm
% r
r
c
C
r
30
PortMic
1 mm
r
c
%
r
Penz
5 mm
r
c
r
5 mm
r
c
30
Alm-50
Prp
D Alm-50
Alm-50
20
20
10
Sps 200
200 ppm
ppm
V V Sc
ppm
800
800 HREE
Y Ti
E
Y
Y
HREE
HREE
Ti
Ti
0
F
Grs
G
Grs
H
Grs Pabu
Grs
I
Lage
Grs Trao Kerz
PLi1-5-1
cores rims
Prp
Li
Ti
Penz rims
V V
400
Y
HREE
0
Sc
Li
0
ppm
0
Sc
100
Li
Li
400
Grs Grs
Sc
100
Prp
Sps
10
Sps
Grs
Sps
Prp
Prp
Alm-50 Grs
Prp Sps Sps
cores Portezmic
cores
cores
LC12
PrpSps
rims
Prp Sps
rims
cores
Prp Sps
rims
Sps
Figure 4. Mineral compositions in garnet from monazite-bearing paragneisses and micaschists. (A–D) Garnet zonation profiles from core (c) to rim (r) in almandine (Alm-50), grossular (Grs), pyrope (Prp), and spessartine (Sps) wt %. Corresponding zonations of trace elements, analyzed by LA-ICP-MS, tic marks in bold indicate positions of single craters. See Table 1 for selected analyses. (E–I) Garnet zonation trends in grossular (Grs)–pyrope (Prp)–spessartine (Sps) coordinates. LC12 in sample Pabu marks analysis reported by Jones (1994).
been observed by Jones (1994). Maximum pyrope contents in sample Penz (Prp 12.5%) are slightly lower than in samples from Porz Liogan (13.0–14.6%) and Plage de Portez (13.0–14.3%). The compositions of S2 biotites are always slightly variable and range between XMg 0.45–0.49 (sample Penz) and XMg 0.43–0.48 (sample Portezmic), but a trend toward lower XMg in the north, as described by Jones (1994), is only weakly developed. Garnet-biotite Fe-Mg exchange thermometry and GMBP barometry in such progressively deformed metapelites was performed by analyses from mica and plagioclase corresponding to early (= core) and successively later (= toward rim) stages of the garnet growth. When compositional zoning of garnet in metapelites is a product of continuous reactions involving garnet, mica, plagioclase, staurolite, aluminosilicates, and quartz of lowvariance assemblages (Tracy, 1982), it reflects finite temporal and spatial domains of equilibration. Moreover, when the coexistent
minerals (mica and plagioclase) are preserved within a microstructural domain or occur as inclusions that can be correlated with matrix phases (St-Onge, 1987), their chemical compositions in such “local equilibria” allow evaluation of successive P and T or P-T changes for consecutive stages of garnet growth (Triboulet and Audren, 1985; Audren and Triboulet, 1993; Schulz, 1993). The basic assumption is that domain equilibrium was achieved and preserved within a given volume of rock. A minimum error of ±50 °C and ±1.5 kbar has to be considered. Compositional changes of plagioclase enclosed in garnet and in the matrix can be excluded because of slow volume diffusion rates (Spear et al., 1990). From the slight variations in XMg it can be concluded that biotite behaved as a reservoir. Mica that occurs as inclusion in garnet or in microlithons wrapped by S2 is considered to have grown early and simultaneously with garnet cores. Mica along S2 or in pressure
Cadomian and Variscan metamorphic events in the Léon domain
P (kbar)
Grt-Cpx
Lesneven Gneiss Unit Eclogites (1) Eclogites (2) Lan Kao Kerhorn Pabu Kerz Trao Lage gAm
14 12 10
8 6
c Grt
r GASP
c Grt-Cd
4
M
s-
St+
2
A
400
P (kbar) 14
T (°C) 800
600
Conquet-Penze Unit
PortCo Penz RenS3
PLi3 PLiN1, PLi412 Port, PortMic
10
GMBP
8
lower c
6
upper
c
r
St+
12
gAm
gAm s-
4 M
shadows around garnet documents a later stage of crystallization (Fig. 3A–C). Similarly, early stages of mineral chemical evolution are displayed by plagioclase enclosed in garnets, and/or the cores of zoned plagioclase porphyroblasts. The late stage is documented by the rims of zoned porphyroblasts. Mica and plagioclase of S1i in garnet, early mica in microlithons, and cores of zoned plagioclase were related to early garnet core compositions. Syn-S2 mica and the plagioclase rims coexisted with late Mg-rich garnet rims. Accordingly, P and T for crystallization of Mg-rich garnet rims have been calculated using the garnet-biotite thermometer of Bhattacharya et al. (1992) in combination with the GMBP barometer (Holland and Powell, 1990), as reported above. The Holland and Powell (1990) version of the GASP barometer gives values intermediate between the pressure estimates from biotite-Mg and biotite-Fe in GMBP. Temperatures calculated from garnet rims with the maximum Mg content range between 580 °C (Penz) and 610 °C (PLiN1) in the southern part and ~610 °C in the northern part of the micaschist unit (Fig. 5B); thus, a small difference in temperatures appears. Temperatures calculated from the Mg-poor garnet cores are lower and range between 490 and 570 °C, and apparently are dependent on the exact analysis location in zoned garnets with low Mg contents. Corresponding to decreasing An contents in plagioclase, pressures calculated from garnet rims and cores appear to be systematically ~1 kbar higher in the northern and structurally deeper parts of the section. The maximum pressures are ~8 kbar at Penzer and 9 kbar at Plage de Portez for a temperature of 610 °C (Fig. 5B). The combination of the thermobarometric data from metabasites, paragneisses, and micaschists can be arranged into continuous, single, clockwise P-T paths for each of the lithological units (Fig. 5). A significant increase of metamorphic pressure and temperature is observed toward the lower parts of the crustal pile, which is represented by the Lesneven Gneiss unit. This pattern is in accordance with previous observations (Le Corre et al., 1991; Jones, 1994), suggesting that the central Léon lithotectonic units represent a normal crustal pile, with the highest-grade rocks in the structurally lower part. However, even when hampered by the lack of garnet-bearing rocks in the transition zone, which is best exposed around Le Conquet, the thermobarometric data do not support the existence of a major thrust or metamorphic discontinuity between the gneiss and the micaschist units. As has been stated by Jones (1994), the lithological boundary has a transitional character. It coincides with the occurrence of garnet with flat zoning profiles in Mn. The transitional character of the boundary is further supported by the occurrence of the foliated Pointe des Renards metagranitoid in both units. According to the sparse chronological data (565 ± 40 Ma, Rb-Sr whole rock) and the structural observations (Michot and Deutsch, 1970; Chauris and Hallégouët, 1989), the protolith of the metagranitoid is interpreted to have intruded before the formation of the foliation in the host rocks. However, the present P-T data do not exclude the possibility that the units are both part of a major nappe unit that was underthrusted beneath the Central Armorican Domain, as has been proposed by Rolet et al. (1986, 1994). Numerical thermal
277
2
B
400
600
T (°C) 800
Figure 5. P-T data and P-T paths from the Léon metabasites, paragneisses, and micaschists. (A) Lesneven Gneiss unit. Eclogites (1): P-T data from garnet-clinopyroxene-plagioclase equilibria (Grt-Cpx), involving data presented by Paquette et al., 1987; see text); Eclogites (2): Relative P-T evolution derived from prograde garnet core (c) to rim (r) zonations in eclogites (this study). P-T data from assemblages with green amphibole (gAm) in overprinted eclogites according to the thermobarometer of Zenk and Schulz (2004). P-T data from garnetsillimanite-plagioclase-quartz (GASP) and garnet-cordierite (Grt-Cd) equilibria in paragneisses, calculated with garnet core (c) compositions. Ms–, muscovite-out reaction; St+, staurolite-in reaction. (B) ConquetPenze Micaschist unit. P-T-data from garnet-muscovite-biotite-plagioclase-quartz equilibria in micaschists, with prograde P-T paths derived from garnet core-to-rim (c, r) zonations; note higher pressure in structurally lower parts of the unit. Abbreviations as in panel A. GMPB— garnet-muscovite-biotite-plagioclase geothermobarometer.
278
Schulz et al.
modeling (Davy and Gillet, 1986) has shown that underthrusting during a continental collision could result in different shapes and maximum temperatures of P-T paths in the upper and lower parts of a lithotectonic pile, as is observed from the Léon units. Garnet Trace Element Composition Y and REE distribution between metamorphic accessory phases (e.g., monazite, xenotime) and major phases (e.g., garnet) is likely to be the result of the whole-rock reaction history. Monazite, xenotime, and garnet appear to be the major fractionating phases for Y in metapelites (Pyle and Spear, 1999; Pyle et al., 2001); therefore, trace element zonation in garnet may provide information concerning the relative appearance and chemical composition of monazite. Maximum Y, Sc (50–250 ppm), and V (20–80 ppm) contents in the garnet zonation profiles are dependent on garnet mode and element abundances in the bulk rock. For Y this bulk rock dependency is obvious in the amphibolite-facies micaschist samples with high maximal Y in garnet, as well as for the low-Y garnets in the high-grade aluminous paragneisses. The maximum Li contents in garnet are systematically lower (8–74 ppm) in the aluminous paragneisses and high-grade rocks than in the micaschists (Li 26–126 ppm), where the garnet is associated with staurolite rich in Li (500–1000 ppm). Heavy REE (HREE) and Y display a strictly linear correlation in garnet (Fig. 4A–D). Like Mn, both Y and HREE should decrease from garnet core to rim at increasing temperature when garnet appears as the only Mn, Y, and HREE fractionating phase in equilibrium with xenotime and monazite, which are consumed (Pyle and Spear, 2000). At low temperatures, xenotime is supposed to be a stable phase (Pyle et al., 2001). When the mode or volume of garnet in a xenotime-bearing rock successively increases during metamorphism, garnet, which crystallizes at higher grades, should be progressively lower in Y and HREE. This property could explain the low Y and HREE in garnets of aluminous paragneiss samples where garnet occurs at high modes (10–15 vol%) and bulk rock Y shows a broad variation of 22–35 ppm, as in the other metasediments. Garnet in the Conquet-Penze micaschists, which crystallized at lower temperatures and recorded a prograde P-T path, is rich in both Y (up to 1000 ppm) as well as HREE (up to 500 ppm) when compared to garnet in the aluminous paragneisses. Furthermore, decreases in Y and HREE are observed in the garnet rims of samples Penz, PLiN1, and PortMic (Fig. 4C and D, Table 1). However, in samples Penz and PortMic, the Y and REE are very low in the Mn-rich cores of garnet that crystallized at low temperatures. The Y contents strongly increase toward the garnet inner rims. This trend may indicate that the Y was initially bound in xenotime and/or monazite and then liberated by the breakdown of these phases under prograde metamorphic conditions (Fig. 4C and D). MINERAL CHEMISTRY AND AGE OF MONAZITE Bulk rock REE absolute abundances and pattern of monazite-bearing metasediments match the post-Archean average
Australian sedimentary rock (PAAS) composition reported in Rollinson (1993). This match indicates that the occurrence of monazite is not related to any specific concentration of REE in the metasediments. The bulk rock REE contents increase, while MgO and Y contents (20–40 ppm) decrease, regardless of metamorphic grade and abundance or absence of garnet. Al2O3 is >20 wt% and K2O ~1.0 wt% in aluminous paragneisses compared to Al2O3 of 15–17 wt% and K2O of 3–4.5 wt% in the other metasediments. The CaO of monazite-bearing samples is comparably low (0.5– 2.7 wt%), supporting the general observation that monazite is not stable when calcsilicate phases occur (Finger et al., 1998). Monazite grain sizes increase with metamorphic grade and range from 20–100 μm in amphibolite-facies micaschists up to 250 μm in high-grade aluminous paragneisses. Two distinct age populations of monazites, Variscan and Cadomian, as detailed below, have been found in the Léon metasediments (Table 2). The observation of two separated monazite age groups is supported by mineral chemical data. Compositions of natural monazite vary in LREE, Th, Y, Ca, and Si (Franz et al., 1996). Trivalent Y substitutes for LREE, and Th4+ and U4+ can be substituted by Th or U + Si = REE + P (huttonite substitution), with Si replacing P in the tetrahedral site, and by Th + U + Ca = 2REE (brabantite substitution), with Ca replacing in addition REE in the 8-fold site (Spear and Pyle, 2002). In the studied monazites, an excellent linear correlation between Th + U and Ca exists, suggesting that the brabantite exchange is the dominant substitution (Fig. 6A). Furthermore, Th and U display broad variations in both Variscan and Cadomian monazites, as is obvious from the U/Pb-Th/Pb diagram (Fig. 7H), which signalizes an unlimited substitution. The REE roughly correlate negatively and linearly with Y2O3 (Fig. 6B). Significant variations of REE and Y in monazite with metamorphic grade were reported by Heinrich et al. (1997), Pyle et al. (2001), Finger et al. (2002), and Spear and Pyle (2002). Therefore, it is interesting to compare compositional parameters of Variscan and Cadomian monazites. The Y2O3 in Cadomian monazite is elevated at ~2.2 wt%. In contrast, considerable variation of Y2O3 is observed from Variscan monazite: high-grade aluminous paragneisses contain monazite low in Y2O3 (~0.2 wt%), intermediate Y2O3 (1.0–1.8 wt%) contents are observed in monazite from amphibolite-facies garnet-bearing rocks, and high Y2O3 contents of >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
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rn
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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
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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
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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 (<4 kbar and 700 °C) and the lack (or rarity) of garnet (Weber et al., 1985; Ballèvre et al., 2001). The boundaries of the Léon are masked and overprinted by the Carboniferous Morlaix basin, the Variscan granites, and the dextral Variscan North Armorican shear zone and their related displacement zones. However, in the light of the Cadomian monazite ages, the Léon domain appears to be a part of the external Cadomian belt, as exposed in the St. Malo or Fougères units and thus is not entirely exotic within the North Armorican domain. In conclusion, in the Léon domain several features typical of the Saxo-Thuringian zone of the Variscan belt, as defined by Kossmat (1927), can be recognized. One characteristic of the Saxo-Thuringian zone is the occurrence of Cadomian remnants, displayed by the Cadomian monazite ages. According to their bulk rock chemistry, the Léon metasediments were deposited in an active continental margin setting as a part of the Briovérien sequences. Such Cadomian elements still have to be better constrained, for instance, by radiometric dating of the Pointe des
Renards metagranitoid and other intrusive rocks. Other features of the Saxo-Thuringian zone observed in the Léon are Variscan granitic plutonism, deformation, and metamorphism to various degrees. The eclogite-facies metamorphism has been taken as an argument for a South Armorican affinity. However, apart from different bulk rock characteristics, the P-T paths and the tectonic relationships of the South Armorican high-pressure metabasites in Baie d´Audierne (Lucks et al., 2002), Ile de Groix (Schulz et al., 2001), and Champtoceaux (Ballèvre et al., 1994) considerably differ from the marked decompression-heating path of the Léon eclogites, regardless of the ages of the high-pressure events, which still need to be made more precise. To the southeast and east, Early Paleozoic to Devonian and Carboniferous sedimentary records are preserved. The central Léon metamorphic units represent the footwall of these sedimentary basins. As regards the Upper Devonian “phase bretonne,” known from the stratigraphic record (Rolet et al., 1986, 1994), the metamorphic monazite ages postdate this event. However, the S1 and S2 deformation structures, with associated crystallization of garnet, aluminosilicates, and staurolite, predate the Carboniferous intrusion of the Saint Renan–Kersaint granite (340–330 Ma). Then the range of Variscan monazite ages (sample PLiN1, 340 ± 16 Ma; sample Port, 304 ± 21 Ma) link this event to a Late Carboniferous stage, with overprinting of the planar structures by dextral shearing. The younger Variscan monazite ages indicate that the Léon units were exhumed no earlier than the Upper Carboniferous, presumably in line with the dextral wrench tectonics. The P-T data and P-T paths in combination with the monazite ages emphasize that the central Léon units underwent a common structural and metamorphic evolution during the Variscan orogeny. Inverted metamorphic gradients or superposition of high-pressure rocks on lower-pressure units, which provide arguments for nappe tectonics in other parts of the Variscan belt (e.g., Matte, 1991), are not evident in the Léon. This observation does not exclude the possibility that the units may represent allochthonous or parautochthonous crust. Rolet et al. (1986, 1994) proposed that the Léon units were underthrusted toward the southeast or east beneath the Central Armorican domain during a Variscan collision. In this model, a South Armorican provenance of the nappes is not favorable, despite the possible re-arrangement of tectonic domains during Carboniferous dextral wrench tectonics. As supported by the Cadomian monazite ages, an origin of the Léon units as former parts of a Cadomian realm, situated to the northwest or west appears to be more likely. When interpreted in the light of geodynamic reconstructions of the western part of the Variscan belt (Matte, 2002; Shelley and Bossière, 2002; Stampfli et al., 2002; Cartier and Faure, 2004), our findings do not serve as an additional argument for the existence of an Early Paleozoic oceanic realm (Medio-European Ocean, Massif Central Ocean, or South Armorican Ocean) and suture zone between the southern margin of an Armorican microplate and Gondwana. They rather support the notion that the Léon units were parts of a suture zone along the northern boundary of the Armorican microplate, hence related to the margin of a former Rheic Ocean between Armorica and Avalonia.
Cadomian and Variscan metamorphic events in the Léon domain ACKNOWLEDGMENTS We commemorate the Breton geologist Claude Audren from Géosciences Rennes, who guided the first author to the Léon in 2001. The constructive reviews by J.F. von Raumer, Fribourg, and J. Rolet, Plouzane, are gratefully acknowledged. We thank U. Kroner, Freiberg, for giving us insight to his manuscript. Microprobe analyses were facilitated by H.-P. Meyer at Mineralogisches Institut Universität Heidelberg and by U. Schüßler, Würzburg. Technical support was provided by R. Baur, P. Spaethe, and K.-P. Kelber at the Institut für Mineralogie in Würzburg. The project was financed through the Deutsche Forschungsgemeinschaft (SCHU 676/10). REFERENCES CITED Ashworth, J.R., and Evirgen, M.M., 1985, Plagioclase relations in pelites, central Menderes Massif, Turkey. II. Perturbation of garnet-plagioclase geobarometers: Journal of Metamorphic Geology, v. 3, p. 219–229. Audren, C., and Triboulet, C., 1993, P-T-t-deformation paths recorded by kinzigites during diapirism in the western Variscan belt (Golfe du Morbihan, southern Brittany, France): Journal of Metamorphic Geology, v. 11, p. 337–356. Balé, P., and Brun, J.-P., 1986, Les complexes métamorphiques du Léon (NW Bretagne): un segment du domaine éo-hercynien sud armoricain translaté au Dévonien: Bulletin Société Géologique de France, v. 1986, série 3, p. 471–477. Ballèvre, M., Marchand, J., Godard, G., Goujou, J.C., and Wyns, R., 1994, Eo-Hercynian events in the Armorican Massif, in Keppie, J.D., ed., Pre-Mesozoic geology in France and related areas: Berlin, Springer-Verlag, p. 183–194. Ballèvre, M., Le Goff, E., and Hébert, R., 2001, The tectonothermal evolution of the Cadomian belt of northern Brittany, France: A Neoproterozoic volcanic arc: Tectonophysics, v. 331, p. 19–43, doi: 10.1016/S0040-1951(00)00234-1. Bhatia, M.R., and Crook, K.A.W., 1986, Trace element characteristics of greywackes and tectonic setting discrimination of sedimentary basins: Contributions to Mineralogy and Petrology, v. 92, p. 181–193, doi: 10.1007/BF00375292. Bhattacharya, A., Mohanty, L., Maji, A., Sen, S.K., and Raith, M., 1992, Nonideal mixing in the phlogopite-annite binary: Constraints from experimental data on Fe-Mg partitioning and a reformulation of the garnet-biotite geothermometer: Contributions to Mineralogy and Petrology, v. 111, p. 87–93, doi: 10.1007/BF00296580. Brätz, H., and Klemd, R., 2002, Analysis of rare earth elements in geological samples by laser ablation–inductively coupled plasma mass spectrometry (LA-ICP-MS): Online publication 5988–6305EN, Agilent Technologies. Brown, M., and Dallmeyer, R.D., 1996, Rapid Variscan exhumation and the role of magma in core complex formation: Southern Brittany metamorphic belt, France: Journal of Metamorphic Geology, v. 14, p. 361–379, doi: 10.1111/j.1525-1314.1996.00361.x. Brun, J.-P., and Balé, P., 1990, Cadomian tectonics in northern Brittany, in D´Lemos, R.S., Strachan, R.A., and Topley, C.G., eds., The Cadomian orogeny: London, Geological Society of London Special Publication 51, p. 95–114. Brun, J.-P., Guennoc, P., Truffert, C., Vairon, J., and The ARMOR Working Group of the GeoFrance 3-D program, 2001, Cadomian tectonics in northern Brittany: A contribution of 3-D crustal-scale modelling: Tectonophysics, v. 331, p. 229–246. Cabanis, B., and Godard, G., 1987, Les éclogites du pays de Léon (Nord-Ouest du Massif armoricain): étude pétrologique et géochimique; implications géodynamiques: Bulletin Société Géologique de France, v. 1987, série 8 III, no. 6, p. 1133–1142. Cabanis, B., Michot, J., and Deutsch, S., 1977, Remise en question de la datation géochronologique des gneiss de Brest (Bretagne occidentale): Comptes Rendus de l´Académie des Sciences Paris, v. 284, série D, p. 883–886. Cabanis, B., Chantraine, J., and Herrouin, Y., 1979a, Le Bassin de Morlaix, unité circonscrite et indépendante dans le contexte géologique régional: Bulletin du Bureau de Recherches Géologiques et Minières, v. 4, p. 269–276. Cabanis, B., Peucat, J.-J., Michot, J., and Deutsch, S., 1979b, Remise en cause de l´existence d´un socle orthogneissique antécambrien dans le pays de Léon
283
(domaine nord-armoricain); étude géochronologique par les méthodes Rb/ Sr et U/Pb des orthogneiss de Tréglonou et de Plounevez-Lochrist: Bulletin du Bureau de Recherches Géologiques et Minières, v. 1/4, p. 357–364. Cartier, C., and Faure, M., 2004, The Saint-Georges-sur Loire olisthostrome, a key zone to understand the Gondwana-Armorica boundary in the Variscan belt (Southern Brittany, France): International Journal of Earth Sciences, v. 93, p. 945–958, doi: 10.1007/s00531-004-0398-3. Chantraine, J., Chauris, L., Cabanis, B., Chauris, M.-M., Larsonneur, C., Herrouin, Y., Rabu, D., Lulzac, Y., and Bos, P., 1986, Notice explicative de la feuille Plestin-les-Grèves à 1/50 000: Orléans, Éditions du Bulletin du Bureau de Recherches Géologiques et Minières, 64 p. Chantraine, J., Egal, E., Thiéblemont, D., Le Goff, E., Guerrot, C., Ballèvre, M., and Guennoc, P., 2001, The Cadomian active margin (North Armorican Massif, France): A segment of the North Atlantic Panafrican belt: Tectonophysics, v. 331, p. 1–18, doi: 10.1016/S0040-1951(00)00233-X. Chauris, L., and Hallégouët, B., 1989, Notice explicative Carte géologique de la France (1/50 000), feuille Plouarzel—Ile d´Ouessant (273): Orléans, Éditions du Bulletin du Bureau de Recherches Géologiques et Minières, 69 p. Chauris, L., Marcoux, È., Le Goff, È., Thiéblemont, D., and Carn, A., 1998, Notice explicative Carte géologique de la France (1/50 000), feuille SaintPol-de-Léon (201): Orléans, Éditions du Bulletin du Bureau de Recherches Géologiques et Minières, 167 p. Cocherie, A., and Albarede, F., 2001, An improved U-Th-Pb age calculation for electron microprobe dating of monazite: Geochimica et Cosmochimica Acta, v. 65, p. 4509–4522, doi: 10.1016/S0016-7037(01)00753-0. Cocherie, A., Legendre, O., Peucat, J.J., and Koumelan, A.N., 1998, Geochronology of polygenetic monazites constrained by in-situ microprobe ThU-total lead determination: Implications for lead behaviour in monazite: Geochimica et Cosmochimica Acta, v. 62, p. 2475–2497, doi: 10.1016/ S0016-7037(98)00171-9. Cogné, J., 1988, Precambrian in the Variscides of western France: Armorican Massif and western parts of the Central Massif, in Zoubek, V., ed., Precambrian in younger fold belts: European Variscides, the Carpathians and Balkans: Chichester, John Wiley, p. 409–552. Cogné, J., 1990, The Cadomian orogeny and its influence on the Variscan evolution of western Europe, in D´Lemos, R.S., Strachan, R.A., and Topley, C.G., eds., The Cadomian orogeny: London, Geological Society of London Special Publication 51, p. 305–311. Cogné, J., and Shelley, D., 1966, Structure géologique du secteur des Abers (Nord Finistère): Bulletin Service Carte Géologique Alsace Lorraine, v. 19, no. 1, p. 1–40. Dahl, P.S., Terry, M.P., Jercinovic, M.J., Williams, M.L., Hamilton, M.A., Foland, K.A., Clement, S.M., and Friberg, L.M., 2005, Electron probe (Ultrachron) micronometry of metamorphic monazite: Unraveling the timing of polyphase thermotectonism in the easternmost Wyoming craton (Black Hills, South Dakota): American Mineralogist, v. 90, p. 1712–1728, doi: 10.2138/am.2005.2002. Dallmeyer, R.D., Brown, M., D´Lemos, R.S., and Strachan, R.A., 1993, Variable Variscan thermal rejuvenation in the St. Malo region, Cadomian orogen, France: Evidence from 40Ar/39Ar mineral ages: Journal of Metamorphic Geology, v. 11, p. 137–154. Dalziel, I.W.D., 1997, Neoproterozoic-Palaeozoic geography and tectonics: Review, hypothesis, environmental speculation: Geological Society of America Bulletin, v. 109, p. 16–42, doi: 10.1130/0016-7606(1997)109<0016: ONPGAT>2.3.CO;2. Davy, P., and Gillet, P., 1986, The stacking of thrust slices in collision zones and its thermal consequences: Tectonics, v. 5, p. 913–929. Deutsch, S., and Chauris, L., 1965, Âge des quelques formations cristallophylliennes et granitiques du pays de Léon, Finistère: Comptes Rendus de l´Académie des Sciences Paris, v. 260, p. 615–617. Egal, E., Guerrot, C., Le Goff, E., Thiéblemont, D., and Chantraine, J., 1996, The Cadomian orogeny revisited in northern Brittany (France), in Nance, R.D., and Thompson, M.D., eds., Avalonian and related Peri-Gondwanan terranes in the circum-North Atlantic: Boulder, Colorado, Geological Society of America Special Paper 304, p. 281–317. Ellis, D.J., and Green, D.H., 1979, An experimental study of the effect of Ca upon garnet-clinopyroxene Fe-Mg exchange equilibria: Contributions to Mineralogy and Petrology, v. 71, p. 13–22, doi: 10.1007/BF00371878. Finger, F., and Helmy, H.M., 1998, Composition and total-Pb model ages of monazite from high-grade paragneisses in the Abu Swayel area, southern Eastern Desert, Egypt: Mineralogy and Petrology, v. 62, p. 269–289, doi: 10.1007/BF01178032.
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Schulz et al.
Finger, F., Broska, I., Roberts, M.P., and Schermaier, A., 1998, Replacement of primary monazite by apatite-allanite-eoidote coronas in an amphibolitefacies granite gneiss from the eastern Alps: American Mineralogist, v. 83, p. 248–258. Finger, F., Krenn, E., Riegler, G., Romano, S., and Zulauf, G., 2002, Resolving Cambrian, Carboniferous, Permian and Alpine monazite generations in the polymetamorphic basement of eastern Crete (Greece) by means of the electron microprobe: Terra Nova, v. 14, p. 233–240, doi: 10.1046/j.13653121.2002.00415.x. Floyd, P.A., Winchester, J.A., and Park, R.G., 1989, Geochemistry and tectonic setting of Lewisian clastic metasediments from the Early Proterozoic Loch Maree Group of Gairloch, NW Scotland: Precambrian Research, v. 45, p. 203–214, doi: 10.1016/0301-9268(89)90040-5. Foster, G., Kinny, P., Vance, D., Prince, C., and Harris, N., 2000, The significance of monazite U-Th-Pb age data in metamorphic assemblages; A combined study of monazite and garnet chronometry: Earth and Planetetary Science Letters, v. 181, p. 327–340, doi: 10.1016/S0012-821X(00)00212-0. Foster, G., Gibson, H.D., Parrish, R., Horstwood, M., Fraser, J., and Tindle, A., 2002, Textural, chemical and isotopic insights into the nature and behaviour of metamorphic monazite: Chemical Geology, v. 191, p. 183–207, doi: 10.1016/S0009-2541(02)00156-0. Franke, W., 1989, Tectonostratigraphic units in the Variscan belt of central Europe, in Dallmeyer, R.D., ed., Terranes in the circum-Atlantic Paleozoic orogens: Boulder, Colorado, Geological Society of America Special Paper 230, p. 67–90. Franz, G., Andrehs, G., and Rhede, D., 1996, Crystal chemistry of monazite and xenotime from Saxothuringian-Moldanubian metapelites, NE Bavaria, Germany: European Journal of Mineralogy, v. 8, p. 1097–1118. Friedl, G., 1997, U/Pb-Datierungen an Zirkonen und Monaziten aus Gesteinen vom österreichischen Anteil der Böhmischen Masse [Ph.D. thesis]: Salzburg, Universität Salzburg, 242 p. Ganguly, J., and Saxena, S.K., 1984, Mixing properties of aluminosilicate garnets: Constraints from natural and experimental data, and applications to geothermo-barometry: American Mineralogist, v. 69, p. 88–97. Ganguly, J., Cheng, W., and Tirone, M., 1996, Thermodynamics of aluminosilicate garnet solid solution: New experimental data, an optimized model, and thermometric applications: Contributions to Mineralogy and Petrology, v. 126, p. 137–151, doi: 10.1007/s004100050240. Gerya, T.V., Perchuk, L.L., Triboulet, C., Audren, C., and Sez’ko, A.I., 1997, Petrology of the Tumanshet zonal metamorphic complex, Eastern Sayan: Petrology (Moscow), v. 5, p. 503–533. Godard, G., and Mabit, J.-L., 1998, Peraluminous sapphirine formed during retrogression of a kyanite-bearing eclogite from Pays de Léon, Armorican Massif, France: Lithos, v. 43, p. 15–29, doi: 10.1016/S00244937(98)00004-8. Goré, B., and Le Corre, C., 1987, Cinématique hercynienne du cisaillement nord-armoricain à la bordure du granite syntectonique de Saint RenanKersaint (Finistère): Bulletin Société Géologique de la France, v. 1987, série 8, no. 5, p. 811–819. Govindaraju, K., 1994, Compilation of working values and sample description for 383 geostandards: Geostandards Newsletters, v. 18, special issue, p. 1–158. Hébert, R., 1995, Evidence for multiple high-T metamorphism and regional migmatization within the Cadomian belt of northern Brittany, France: Journal of the Geological Society of London, v. 152, p. 213–216. Heinrich, W., Andrehs, G., and Franz, G., 1997, Monazite-xenotime miscibility gap thermometry. I. An empirical calibration: Journal of Metamorphic Geology, v. 15, p. 3–16, doi: 10.1111/j.1525-1314.1997.t01-1-00052.x. Holland, T.J.B., 1983, The experimental determination of activities in disordered and short-range ordered jadeitic pyroxene: Contributions to Mineralogy and Petrology, v. 82, p. 214–220, doi: 10.1007/BF01166616. Holland, T.J.B., and Powell, R., 1990, An enlarged and updated internally consistent thermodynamic dataset with uncertainties and correlations: The system K2O-Na2O-CaO-MgO-MnO-FeO-Fe2O3-Al2O3-TiO2-SiO2-C-H2O2: Journal of Metamorphic Geology, v. 8, p. 89–124. Hollister, L.S., 1966, Garnet zoning: An interpretation based on the Rayleigh fractionation model: Science, v. 154, p. 1647–1651, doi: 10.1126/ science.154.3757.1647. Horn, I., Hinton, R.W., Jackson, S.E., and Longerich, H.P., 1997, Ultra-trace element analysis of NIST SRM 616 and 614 using laser ablation microprobe– inductively coupled plasma–mass spectrometry (LAM-ICP-MS): A comparison with secondary ion mass spectrometry (SIMS): Geostandard Newsletters, v. 21, p. 191–203.
Jackson, S.E., Longerich, H.P., Dunning, G.R., and Fryer, B.J., 1992, The application of laser ablation microprobe–inductively coupled plasma–mass spectrometry (LAM-ICP-MS) to in situ trace element determination in minerals: Canadian Mineralogist, v. 30, p. 1049–1064. Jones, K.A., 1993, Phase relations in Al2SiO5 polymorphs; Le Conquet region, north-western Brittany, France: Proceedings of the Ussher Society, v. 8, p. 138–144. Jones, K.A., 1994, Progressive metamorphism in a crustal-scale shear zone: An example from the Léon region, north-west Brittany, France: Journal of Metamorphic Geology, v. 12, p. 69–88. Jones, K.A., and Brown, M., 1990, High-temperature “clockwise” P-T paths and melting in the development of regional migmatites: An example from southern Brittany, France: Journal of Metamorphic Geology, v. 8, p. 551–578. Kossmat, F., 1927, Die Gliederung des varistischen Gebirges: Abhandlungen Sächsisches Geologisches Landesamt, v. 1, p. 1–39. Krogh, E.J., 1988, The garnet-clinopyroxene Fe-Mg geothermometer—A reinterpretation of existing experimental data: Contributions to Mineralogy and Petrology, v. 99, p. 44–48, doi: 10.1007/BF00399364. Leake, B.E., Woolley, A.R., Arps, C.E.S., Birch, W.D., Gilbert, M.C., Grice, J.D., Hawthorne, F.C., Kato, A., Kisch, H.J., Krichovichev, V.G., Linthout, K., Laird, J., Mandarino, J.A., Maresch, W.V., Nickel, E.H., Rock, N.M.S., Schumacher, J.C., Smith, D.C., Stephenson, N.C.N., Ungaretti, L., Whittaker, E.J.W., and Youzhi, G., 1997, Nomenclature of amphiboles: Report of the Subcommittee on Amphiboles of the International Mineralogical Association, Commission on New Minerals and Mineral Names: American Mineralogist, v. 82, p. 1019–1037. Le Corre, C., Bale, P., and Georget, Y., 1989, Le Léon: un domaine exotique au Nord-Ouest de la chaine varisque armoricaine (France): Geodinamica Acta, v. 3, no. 2, p. 57–71. Le Corre, C., Auvray, B., Ballèvre, M., and Robardet, M., 1991, Le Massif Armoricain: Sciences Géologiques Bulletin Strasbourg, v. 44, no. 1–2, p. 31–103. Leutwein, F., Chauris, L., Sonet, J., and Zimmermann, J.-L., 1969, Etudes géochroniques et géotectoniques dans le Nord-Finistère: Bulletin Sciences de la Terre, v. 14/44, p. 331–358. Linnemann, U., and Romer, R.L., 2002, The Cadomian orogeny in SaxoThuringia, Germany: Geochemical and Nd-Sr-Pb isotopic characterization of marginal basins with constraints to geotectonic setting and provenance: Tectonophysics, v. 352, p. 33–64, doi: 10.1016/S0040-1951(02)00188-9. Longerich, H.P., Jackson, S.E., Fryer, B.J., and Strong, D.F., 1993, The laser ablation microprobe–inductively coupled plasma–mass spectrometer (LAMICP-MS): Geoscience Canada, v. 20, p. 21–27. Lucks, H., Schulz, B., Audren, C., and Triboulet, C., 2002, Variscan P-T evolution of garnet pyroxenites and amphibolites in the Baie d’Audierne metamorphic series, southern Brittany (France), in Martinez-Catalán, J.R., Hatcher, R.D., Arenas, R., and Diaz Garcia, F., eds., Variscan-Appalachian dynamics: The building of the Upper Palaeozoic basement: Boulder, Colorado, Geological Society of America Special Paper 364, p. 89–103. Ludwig, K.R., 2001, Users manual for Isoplot/Ex rev. 2.49. A geochronological toolkit for Microsoft Excel: Berkeley, California, Berkeley Geochronology Center Special Publication 1a, 55 p. Martinez-Catalan, J.R., 1990, A non-cylindrical model for the north-west Iberian allochthonous terranes and their equivalents in the Hercynian belt of Western Europe: Tectonophysics, v. 179, p. 253–272, doi: 10.1016/00401951(90)90293-H. Matte, P., 1991, Accretionary history and crustal evolution of the Variscan belt in Western Europe: Tectonophysics, v. 196, p. 309–337, doi: 10.1016/00401951(91)90328-P. Matte, P., 2002, Variscides between the Appalachian and the Urals: Similarities and differences between Paleozoic subduction and collision belts, in Martinez-Catalán, J.R., Hatcher, R.D., Arenas, R., and Diaz Garcia, F., eds., Variscan-Appalachian dynamics: The building of the Upper Palaeozoic basement: Boulder, Colorado, Geological Society of America Special Paper 364, p. 239–251. Michot, J., and Deutsch, S., 1970, U/Pb zircon ages and polycyclism of the Gneiss de Brest and the adjacent formations (Brittany): Eclogae geologicae Helveticae, v. 63/1, p. 215–227. Montel, J.-M., Veschambre, M., and Nicollet, C., 1994, Datation de la monazite à la microsonde électronique: Comptes Rendus de l´Académie des Sciences Paris, série II, v. 318, p. 1489–1495. Montel, J.-M., Foret, S., Veschambre, M., Nicollet, C., and Provost, A., 1996, A fast, reliable, inexpensive in-situ dating technique: Electron microprobe
Cadomian and Variscan metamorphic events in the Léon domain ages on monazite: Chemical Geology, v. 131, p. 37–53, doi: 10.1016/00092541(96)00024-1. Outin, J.-M., Thomas, É., Hallégouët, B., Queté, Y., Le Goffic, M., Rivière, J.-M., and Ballèvre, M., 2000, Note explicative, Carte géologique de la France (1/50 000), feuille Plabennec (238): Orléans, Èditions du Bulletin du Bureau de Recherches Géologiques et Minières, 109 p. Paquette, J.-L., Balé, P., Ballèvre, M., and Georget, Y., 1987, Géochronologie et géochimie des éclogites du Léon: nouvelles contraintes sur l´évolution géodynamique du Nord-Ouest du Massif Armoricain: Bulletin de Minéralogie, v. 110, p. 683–696. Paris, F., and Robardet, M., 1994, Paleogeographic synthesis, in Keppie, J.D., ed., Pre-Mesozoic geology in France and related areas: Berlin, Heidelberg, Springer-Verlag, p. 172–176. Passchier, C.W., and Trouw, R.A.J., 1996, Microtectonics: Berlin, Heidelberg, Springer-Verlag, 289 p. Pearce, N.J.G., Perkins, W.T., Westgate, J.A., Gorton, M.P., Jackson, S.E., Neal, C.R., and Chenery, S.P., 1997, A compilation of new and published major and trace element data for NIST SRM 610 and NIST SRM 612 glass reference materials: Geostandards Newsletters, v. 21, p. 115–144. Perchuk, L.L., 1991, Derivation of a thermodynamically consistent set of geothermometers and barometers for metamorphic and magmatic rocks, in Perchuk, L.L., ed., Progress in metamorphic and magmatic petrology: Cambridge, Cambridge University Press, p. 93–111. Peucat, J.J., Auvray, B., Hirbec, Y., and Calvez, J.Y., 1984, Granites et cisaillements hercyniens dans le nord du Massif armoricain: géochronologie RbSr: Bulletin de la Société Géologique de France, v. 6, p. 1365–1373. Plyusnina, L.P., 1982, Geothermometry and geobarometry of plagioclase-hornblende bearing assemblages: Contributions to Mineralogy and Petrology, v. 80, p. 140–146, doi: 10.1007/BF00374891. Powell, R., and Holland, T.J.B., 1993, On the formulation of simple mixing models for complex phases: American Mineralogist, v. 78, p. 1174–1180. Pyle, J.M., and Spear, F.S., 1999, Yttrium zoning in garnet: Coupling of major and accessory phases during metamorphic reactions: Geological Materials Research, v. 1, no. 6, p. 1–49. Pyle, J.M., and Spear, F.S., 2000, An empirical garnet (YAG)-xenotime thermometer: Contributions to Mineralogy and Petrology, v. 138, p. 51–58, doi: 10.1007/PL00007662. Pyle, J.M., and Spear, F.S., 2003, Four generations of accessory-phase growth in low-pressure migmatites from SW New Hampshire: American Mineralogist, v. 88, p. 338–351. Pyle, J.M., Spear, F.S., Rudnick, R.L., and McDonough, W.F., 2001, Monazitexenotime-garnet equilibrium in metapelites and a new monazite-garnet thermometer: Journal of Petrology, v. 42, p. 2083–2107, doi: 10.1093/ petrology/42.11.2083. Robardet, M., 2002, Alternative approach to the Variscan belt in southwestern Europe: Preorogenic paleobiogeographical constraints, in Martinez-Catalán, J.R., Hatcher, R.D., Arenas, R., and Diaz Garcia, F., eds., Variscan-Appalachian dynamics: The building of the Upper Palaeozoic basement: Boulder, Colorado, Geological Society of America Special Paper 364, p. 1–16. Rolet, J., 1994, Introduction to structure and metamorphism in the Armorican Massif, France, in Keppie, J.D., ed., Pre-Mesozoic geology in France and related areas: Berlin, Heidelberg, Springer-Verlag, p. 177–178. Rolet, J., Le Gall, B., Darboux, J.-R., Thonon, P., and Gravelle, M., 1986, L’évolution géodynamique dévono-carbonifère de l’extrémité occidentale de la chaîne hercynienne d’Europe sur le transect Armorique–Cornwall: Bulletin Société Géologique de France, v. 1986, série 8/1, p. 43–54. Rolet, J., Gresselin, F., Jegouzo, P., Ledru, P., and Wyns, R., 1994, Intracontinental Hercynian events in the Armorican Massif, in Keppie, J.D., ed., Pre-Mesozoic geology in France and related areas: Berlin, Heidelberg, Springer-Verlag, p. 195–219. Rollinson, H.R., 1993, Using geochemical data: Evaluation, presentation, interpretation: London, Longman Group, 352 p. Roser, B.P., and Korsch, R.J., 1986, Determination of tectonic setting of sandstone-mudstone suites using SiO2 content and K2O/Na2O-ratio: Journal of Geology, v. 94, p. 635–651. Schulz, B., 1993, P-T-deformation paths of Variscan metamorphism in the Austroalpine basement: Controls on geothermobarometry from microstructures in progressively deformed metapelites: Swiss Bulletin of Mineralogy and Petrology, v. 73, p. 257–274. Schulz, B., Audren, C., and Triboulet, C., 1998, Regional vs contact metamorphism in the vicinity of late-Variscan granites (Central Armorican Domain, Brittany, France): International Journal of Earth Sciences, v. 87, p. 78–93.
285
Schulz, B., Triboulet, C., Audren, C., Gilg, A., and Pfeifer, H.-R., 2001, Two-stage prograde and retrograde Variscan metamorphism of glaucophane-eclogites, blueschists and greenschists from Ile de Groix, Brittany: International Journal of Earth Sciences, v. 90, p. 871–889, doi: 10.1007/s005310000156. Shelley, D., and Bossière, G., 2000, A new model for the Hercynian orogen of Gondwanan France and Iberia: Journal of Structural Geology, v. 22, p. 757–776, doi: 10.1016/S0191-8141(00)00007-9. Shelley, D., and Bossière, G., 2002, Megadisplacements and the Hercynian orogen of Gondwanan France and Iberia, in Martinez-Catalán, J.R., Hatcher, R.D., Arenas, R., and Diaz Garcia, F., eds., Variscan-Appalachian dynamics: The building of the Upper Palaeozoic basement: Boulder, Colorado, Geological Society of America Special Paper 364, p. 209–222. Spear, F.S., 1993, Metamorphic phase equilibria and pressure-temperature-time paths: Washington, D.C., Mineralogical Society of America Monograph Series 1, 799 p. Spear, F.S., and Pyle, J.M., 2002, Apatite, monazite and xenotime in metamorphic rocks, in Kohn, M., Rakovan, J., and Hughes, J.M., eds., Phosphates—Geochemical, geobiological and materials importance: Reviews in Mineralogy and Geochemistry, v. 48, p. 293–335. Spear, F.S., Kohn, M.J., Florence, F.P., and Menard, T., 1990, A model for garnet and plagioclase growth in pelitic schists: Implications for thermobarometry and P-T path determinations: Journal of Metamorphic Geology, v. 8, p. 683–696. Stampfli, G.M., von Raumer, J.F., and Borel, G.D., 2002, Paleozoic evolution of pre-Variscan terranes: From Gondwana to the Variscan collision, in Martinez-Catalán, J.R., Hatcher, R.D., Arenas, R., and Diaz Garcia, F., eds., Variscan-Appalachian dynamics: The building of the Upper Palaeozoic basement: Boulder, Colorado, Geological Society of America Special Paper 364, p. 263–280. St-Onge, M.R., 1987, Zoned poikiloblastic garnets: P-T paths and synmetamorphic uplift through 30 km of structural depth, Wopmay orogen, Canada: Journal of Petrology, v. 28, p. 1–21. Strachan, R.A., Treloar, P.J., Brown, M., and D’Lemos, R.S., 1989, Cadomian terrane tectonics and magmatism in the Armorican Massif: Journal of the Geological Society of London, v. 146, p. 423–426. Suzuki, K., Adachi, M., and Kajizuka, I., 1994, Electron microprobe observations of Pb diffusion in metamorphosed detrital monazites: Earth and Planetary Science Letters, v. 128, p. 391–405, doi: 10.1016/0012821X(94)90158-9. Sylvester, P.J., 2003, Trace element analysis of fused whole rock glasses by laser ablation ICPMS, in Sylvester, P.J., ed., Laser-ablation ICPMS in the earth sciences. Principles and applications: Mineralogical Association of Canada Short Course Series 29, p. 147–162. Tracy, R.J., 1982, Compositional zoning and inclusions in metamorphic minerals, in Ferry, J.M., ed., Characterization of metamorphism through mineral equilibria: Washington, D.C., Mineralogical Society of America Reviews in Mineralogy 10, p. 355–397. Triboulet, C., 1992, The (Na-Ca)amphibole-albite-chlorite-epidote-quartz geothermobarometer in the system S-A-F-M-C-N-H2O. 1. An empirical calibration: Journal of Metamorphic Geology, v. 10, p. 545–556. Triboulet, C., and Audren, C., 1985, Continuous reactions between biotite, garnet, staurolite, kyanite-sillimanite-andalusite and P-T-time-deformation path in micaschists from the estuary of the River Vilaine, south Brittany, France: Journal of Metamorphic Geology, v. 3, p. 91–105. Triboulet, C., and Audren, C., 1988, Controls on P-T-t deformation path from amphibole zonation during progressive metamorphism of basic rocks (estuary of the River Vilaine, south Brittany, France): Journal of Metamorphic Geology, v. 6, p. 117–133. Weber, C., Barbey, P., Cuney, M., and Mertin, H., 1985, Trace element behaviour during migmatization. Evidence for a complex melt-residuum-fluid interaction in the St. Malo migmatic dome (France): Contributions to Mineralogy and Petrology, v. 90, p. 52–62, doi: 10.1007/BF00373041. Williams, M.L., Jercinovic, M.J., and Terry, M.P., 1999, Age mapping and dating of monazite on the electron microprobe; deconvoluting multistage tectonic histories: Geology, v. 27, p. 1023–1026, doi: 10.1130/00917613(1999)027<1023:AMADOM>2.3.CO;2. Zenk, M., and Schulz, B., 2004, Zoned Ca-amphiboles and related P-T evolution in metabasites from the classical Barrovian metamorphic zones in Scotland: Mineralogical Magazine, v. 68, p. 769–786, doi: 10.1180/002 6461046850218. MANUSCRIPT ACCEPTED BY THE SOCIETY 3 OCTOBER 2006 Printed in the USA
Geological Society of America Special Paper 423 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 Gabriel Gutiérrez-Alonso* Departamento de Geología, Facultad de Ciencias, Universidad de Salamanca, 37008 Salamanca, Spain Javier Fernández-Suárez Departamento de Petrología y Geoquímica, Universidad Complutense–Consejo Superior de Investigaciones Científicas, 28040 Madrid, Spain Juan Carlos Gutiérrez-Marco Departamento de Paleontología, Instituto de Geología Económica (Consejo Superior de Investigaciones Científicas– Universidad Complutense de Madrid), 28040 Madrid, Spain Fernando Corfu Institute of Geology, University of Oslo, Blindern, NO-0316 Oslo, Norway J. Brendan Murphy Department of Earth Sciences, St. Francis Xavier University, Antigonish, Nova Scotia, B2G 2W5, Canada Mercedes Suárez Departamento de Geología, Facultad de Ciencias, Universidad de Salamanca, 37008 Salamanca, Spain ABSTRACT Zircons from a >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
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20 cm
0m
Herrería Fm. 1000 m
500 m
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Narcea Shales
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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 <2-μm fraction. Samples were scanned from 2 to 65° 2θ at a 0.05° 2θ/s scan speed. Zircon separation was carried out at the Complutense University of Madrid. The K-bentonite sample was crushed with a jaw crusher and pulverized with a disc mill. Minerals were separated by heavy fraction enrichment on a Wilfley™ table, magnetic separation in a Frantz™ isodynamic separator, and density separation using di-iodomethane (CH2I2). Finally, zircons were handpicked from the nonmagnetic heavy fraction in an alcohol medium under a binocular microscope. The K-bentonite sample contained mainly zircon and pyrite as heavy minerals, which is indicative of limited reworking in the sedimentary environment, in contrast to K-bentonites with the heavy mineral association zircon-tourmaline-rutile, which is a common feature of highly reworked bentonites in which the proportion of remanié zircons is usually high. Zircons were mostly idiomorphic, long prismatic crystals having an aspect ratio of 3:1–4:1, with frequent melt inclusions that in many cases ran along most of the crystal and were thus helpful in the selection of grains devoid of inherited cores. U-Pb analytical work was conducted at the Institute of Geology, University of Oslo. Zircon grains selected for analysis were strongly air-abraded following the method of Krogh (1982). In this case, a ~30–50% size reduction by air-abrasion appears to have completely eliminated lead loss (see below). Abraded zircon grains were washed in 4N HNO3 on a hotplate and rinsed repeatedly with H2O and acetone (with ultrasonication
U-Pb depositional age for the upper Barrios Formation (Armorican Quartzite facies)
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TABLE 1. RESULTS OF ISOTOPE DILUTION–THERMAL IONIZATION MASS SPECTROMETRY U-Pb DATING Apparent age (Ma) 206 † § 206 207 207 206 207 207 †† Sample Wt.* U Th/U Pbcom Pb/ Pb/ Pb/ Pb/ Pb/ Pb/ Pb/ 2σ 2σ 2σ 204 # 238 235 206 238 235 206 (pg) Pb U** U Pb U U Pb (abs) (abs) (μg) (ppm) (abs) AST-1 Z1 5 127 0.14 1.9 1623 0.07703 0.00030 0.6030 0.0035 0.05677 0.00024 478.4 479.1 482.6 Z2 4 152 0.22 2.9 1004 0.07650 0.00031 0.5968 0.0040 0.05658 0.00029 475.2 475.2 475.4 Z3 4 143 0.22 7.3 397 0.07696 0.00032 0.5984 0.0056 0.05640 0.00045 477.9 476.2 468.0 Z4 5 165 0.19 2.9 1675 0.07715 0.00046 0.6010 0.0040 0.05650 0.00034 479.1 477.9 472.0 Z5 4 78 0.29 5.4 298 0.07692 0.00050 0.5999 0.0093 0.05656 0.00081 477.7 477.2 474.6 Z6 5 84 0.30 1.0 1951 0.07701 0.00034 0.5996 0.0034 0.05647 0.00022 478.2 477.0 471.0 *Wt.—Weights no better than 10% when sample weight is ~5 μg. † 208 206 Model Th/U ratio estimated from Pb/ Pb ratio and age of the sample. § Total common Pb in sample, including initial and blank Pb. # Measured ratio, corrected for fractionation and spike contribution. **Corrected for spike, fractionation, blank, and initial common Pb (Stacey and Kramers, 1975). †† 2σ uncertainty calculated by error propagation procedure that takes into account internal measurement statistics and external reproducibility as well as uncertainties in blank and common Pb corrections.
after each rinsing step). After washing and drying, zircon grains were individually weighed in a precision balance. Zircon grains, after the size reduction induced by air-abrasion, weighed between 4 and 5 µg (Table 1). A mixed 205Pb/235U spike was added to the sample after weighing and transferred to the dissolution vessel. Zircon was dissolved in HF (+ HNO3) in Teflon minibombs at ~185 °C for 5 days. Because all zircons weighed less than ~5 μg, no chemical separation of U-Pb was performed, and the whole sample taken in 6N HCl was dried and loaded on outgassed Renium filaments with H3PO4 and silica gel. Isotopic ratios were measured on a Finnigan-MAT™ 262 mass spectrometer by peak jumping on a secondary electron multiplier (ion counting mode). Total procedural blanks were less than 2–5 pg Pb and 0.1–0.3 pg U. The Stacey and Kramers (1975) model was used to subtract initial common Pb in excess of the laboratory blank, and decay constants are those of Jaffey et al. (1971). Concordia plots were created using Isoplot 3.00 (Ludwig, 2003).
RESULTS X-Ray Diffraction In agreement with previously published data, kaolinite is the main mineral component (Fig. 4) of the K-bentonite, and according to the crystalinity index of Hinckley (C.I.H.; Hinckley, 1962), it corresponds to a well-ordered kaolinite (C.I.H. = 0.6). Only some quartz and very small amounts of mica and analcime were detected as impurities in the diffractograms corresponding to powdered samples. These minerals are identified by their principal reflections at 3.34,10, and 5.6 Å, respectively. A very small proportion of pyrophillite (9.2 Å) can be identified in the oriented aggregate (Fig. 4). U-Pb Dating A total of six single-grain analyses were performed on strongly abraded zircons. The analyzed zircon grains (Table 1)
K
K K M P
2
6
10
A 14
M 18
Q Q M
0
10
A
A
20
30
2Ø
40
50
60
Figure 4. XRD patterns of powdered kaolin and oriented aggregate (upper-right corner inset). The y-axis is dimensionless (relative counts); the x-axis is the Bragg angle (2θ). A—analcime; K— kaolinite; M—mica; P—pyrophillite; Q—quartz.
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contain moderate amounts of uranium (78–165 ppm) and a limited range of Th/U ratios (0.15–0.3). The six analyses are concordant and overlap one another at the 2σ confidence level (Fig. 5). The most robust age that can be derived from this data set is the concordia age (Ludwig, 1998). The concordia age algorithm as performed by Isoplot 3.00 (Ludwig, 2003) yields an age of 477.47 ± 0.93 Ma (mean square of weighted deviates and probability of concordance + equivalence are 1.2 and 0.3, respectively, decayconstant errors included; Fig. 5).
0.077
480 0.077
206
0.077
Pb
238
476
U 0.076
DISCUSSION The age of 477.47 ± 0.93 Ma 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 the error of the official age for the Tremadocian-Floian (Lower Ordovician Stage 2) limit at 478.6 ± 1.7 Ma. This figure arises from an interpolation based on the currently accepted radiometric age versus biostratigraphic range timeline for the Ordovician (see Cooper and Sadler, 2004; Sadler and Cooper, 2004). Our age also matches (within error margins) that of related volcanic rocks from northwest Iberia (Loiba volcanics, 475 ± 2 Ma; ValverdeVaquero et al., 2005), Sardinia (Lula porphyroids, 473 ± 13 Ma; Helbing and Tiepolo, 2005), the Saxo-Thuringian basin (Randschiefer beds, 478 ± 1.8 Ma; Schätz et al., 2002) and is very close to other dated K-bentonite layers in surrounding paleogeographic realms, such as north Wales (Compston, 2000). The ca. 477-Ma age slightly postdates extensive plutonic and volcanic activity of the Ollo de Sapo complex (FernándezSuárez et al., 1999; Valverde-Vaquero and Dunning, 2000). This fact, combined with the apparent absence of zircons younger than ca. 550 Ma in beds located a few meters above the K-bentonite in the same section of the Barrios Formation (Fernández-Suárez et al., 2002), suggests that rocks bearing Paleozoic zircons were not eroded in the source area for the Armorican Quartzite facies in this region. An apparent lack of a K-bentonite layer correlatable with the studied one in the southern part of the Cantabrian zone and other realms within the west European Variscan belt can be explained by its erosion during a post-Tremadocian transgression, by the difficulty of preserving ash-fall deposits in high energy or subcontinental environments, or by the failure to recognize such a layer to date. Ash-fall deposits that can be correlated with the dated one in the siliciclastic southwest European Upper Cambrian–Lower Ordovician sequences would provide insights into the nature and episodicity of global volcanic events in the Lower Paleozoic. ACKNOWLEDGMENTS The authors acknowledge the sampling help kindly provided by the staff from Caolines de Merillés S.L. at Mina Conchita (Belmonte de Miranda, Asturias). Helpful reviews by F. Paris and M.J. Bartholomew are also acknowledged. Financial support is from Spanish Ministerio de Educación y Ciencia, Research
0.076
Concordia Age = 477.47 ±0.93 Ma (2σ, decay-const. errs included) MSW (of concordance) = 0.43, Probability (of concordance) = 0.51
472
0.075 0.58
0.59
0.59
0.59 207
0.60
Pb/
235
0.60
0.61
0.61
U
Figure 5. U-Pb concordia plot of the six ID-TIMS single-grain U-Pb analyses (gray ellipses). Ellipses represent 2σ uncertainties. The dark ellipse represents the concordia age and error (Ludwig, 1998, 2003). MSWD—mean square of weighted deviates.
Project CGL2006-00902 (ODRE). This article is a contribution to the International Geological Correlation Program Project 497. REFERENCES CITED Amthor, J.E., Grotzinger, J.P., Schröder, S., Bowring, S.A., Ramezani, J., Martin, M.W., and Matter, A., 2003, Extinction of Cloudina and Namacalathus at the Precambrian-Cambrian boundary in Oman: Geology, v. 31, no. 5, p. 431– 434, doi: 10.1130/0091-7613(2003)031<0431:EOCANA>2.0.CO;2. Aramburu, C., 1989, El Cambro-Ordovícico de la Zona Cantábrica (NO de España) [Ph.D. thesis]: Oviedo, Spain, Universidad de Oviedo, 530 p. Aramburu, C., and García-Ramos, J.C., 1988, Presencia de la discontinuidad Sárdica en la Zona Cantábrica: Geogaceta, v. 5, p. 11–13. Aramburu, C., and García-Ramos, J.C., 1993, La sedimentación Cambro-Ordovícica en la Zona Cantábrica (NO de España): Trabajos de Geología, Oviedo, v. 19, p. 45–73. Aramburu, C., Méndez-Bedia, I., Arbizu, M., and García-López, S., 2004, La secuencia preorogénica, in Vera, J.A., ed., Geología de España: Madrid, Sociedad Geológica de España–Instituto Geológico y Minero de España, p. 27–34. Avigad, D., Kolodner, K., McWilliams, M., Persing, H., and Weissbrod, T., 2003, Origin of northern Gondwana Cambrian sandstone revealed by detrital zircon SHRIMP dating: Geology, v. 31, p. 227–230, doi: 10.1130/00917613(2003)031<0227:OONGCS>2.0.CO;2. Avigad, D., Sandler, A., Kolodner, K., Stern, R.J., McWilliams, M., Miller, N., and Beyth, M., 2005, Mass-production of Cambro-Ordovician quartz-rich sandstone as a consequence of chemical weathering of Pan-African terranes: Environmental implications: Earth and Planetary Science Letters, v. 240, p. 818–826, doi: 10.1016/j.epsl.2005.09.021. Babin, C., 1966, Mollusques bivalves et céphalopodes du Paléozoïque armoricain: Brest, Imprimerie Commerciale et Administrative, 470 p. Babin, C., and Hammann, W., 2001, Une nouvelle espèce de Modiolopsis (Bivalvia) dans l’Arenig (Ordovicien Inférieur) de Daroca (Aragón, Espagne): réflexions sur la denture des bivalves primitifs: Revista Española de Paleontología, v. 16, p. 269–282. Baldwin, C.T., 1977, The stratigraphy and facies associations of trace fossils in some Cambrian and Ordovician rocks from north western Spain, in Crimes, T.P., and Harper, J.C., eds., Trace fossil 2: Geological Journal, Special Issue, v. 9, p. 9–40.
U-Pb depositional age for the upper Barrios Formation (Armorican Quartzite facies) Barrero, O., and Corretgé, L.G., 2002, La diferenciación petrográfica y geoquímica de las rocas ígneas del Arroyo Farandón (Asturias): Geogaceta, v. 32, p. 135–138. Barrois, C., 1882, Recherches sur les terrains anciens des Asturies et de la Galice: Lille, Mémoires de la Société Géologique du Nord, v. 2, 630 p. Barrois, C., 1891, Mémoire sur la Faune du Grès armoricain: Annales de la Société Géologique du Nord, v. 19, p. 134–237. Bassett, M.G., and Cocks, L.R.M., 1974, A review of Silurian brachiopods from Gotland: Fossils and Strata, v. 3, p. 1–56. Bergström, S.M., Löfgren, A., and Maletz, J., 2004, The GSSP of the second (upper) stage of the Lower Ordovician Series: Diabasbrottet at Hunnenberg, province of Västergötland, southwestern Sweden: Episodes, v. 27, p. 265–272. Bergström, S.M., Finney, S.C., Chen, X., Goldman, D., and Leslie, S.A., 2006, Three new Ordovician global stage names: Lethaia, v. 39, p. 287–288. Bonjour, J.L., and Odin, G.-S., 1989, Recherche sur les volcanoclastites des Séries Rouges Initiales en presqu’île de Crozon: premier âge radiométrique de l’Arénig: Géologie de la France, v. 4, p. 3–8. Bonjour, J.L., Peucat, J.J., Chauvel, J.J., Paris, F., and Cornichet, J., 1988, U-Pb zircon dating of the Early Paleozoic (Arenigian) transgression in western Brittany (France): A new constraint for the Lower Paleozoic time-scale: Chemical Geology, v. 72, p. 329–336. Bouyx, E., 1970, Contribution à l’étude des formations Anté-Ordoviciennes de la Meseta Méridionale (Ciudad Real et Badajoz): Memorias del Instituto Geológico y Minero de España, v. 73, p. 1–263. Bowring, S., and Schmitz, M.D., 2003, High-precision U-Pb zircon geochronology and the stratigraphic record: Reviews in Mineralogy and Geochemistry, v. 53, p. 305–326. Brasier, M., Cowie, J., and Taylor, M., 1994, Decision on the PrecambrianCambrian boundary stratotype: Episodes, v. 17, p. 95–100. Brun, J.P., Ballard, J.F., and Le Corre, C., 1991, Identification of Ordovician block tilting in the Hercynian fold belt of central Brittany (France): Field evidence and computer models: Journal of Structural Geology, v. 13, p. 419–429. Cocks, L.R.M., 1993, Triassic pebbles, derived fossils and the Ordovician to Devonian palaeogeography of Europe: Journal of the Geological Society of London, v. 150, p. 219–226. Cocks, L.R.M., 2000, The Early Palaeozoic geography of Europe: Journal of the Geological Society of London, v. 157, p. 1–10. Cocks, L.R.M., 2001, Ordovician and Silurian global geography: Journal of the Geological Society of London, v. 158, p. 197–210. Cocks, L.R.M., and Lockley, M.G., 1981, Reassessment of the Ordovician brachiopods from the Budleigh Salterton pebble bed, Devon: Bulletin of the British Museum of Natural History (Geology), v. 35, p. 111–124. Cocks, L.R.M., and Torsvik, T.H., 2002, Earth geography from 500 to 400 million years ago: A faunal and palaeomagnetic review: Journal of the Geological Society of London, v. 159, p. 631–644. Coke, C., and Gutiérrez-Marco, J.C., 2001, Braquiópodos Linguliformea del Ordovícico Inferior de la Serra do Marão (Zona Centroibérica, N de Portugal): Boletín Geológico y Minero, v. 112, no. 1, p. 33–50. Compston, W., 2000, Interpretations of SHRIMP and isotope dilution zircon ages for the geological time-scale: 1. The early Ordovician and late Cambrian: Mineralogical Magazine, v. 64, p. 43–57, doi: 10.1180/002646100549120. Comte, P., 1937, La série cambrienne et silurienne du León (Espagne): Comptes Rendus de l’Académie des Sciences de Paris, v. 204, p. 604–606. Condon, D., Zhu, M., Bowring, S., Wang, W., Yang, A., and Jin, Y., 2005, U-Pb ages from the Neoproterozoic Doushantuo formation, China: Science, v. 308, p. 95–98, doi: 10.1126/science.1107765. Cooper, R.A., and Sadler, P.M., 2004, The Ordovician Period, in Gradstein, F., Ogg, J., and Smith, A., eds., A geologic time scale 2004: Cambridge, Cambridge University Press, p. 165–187. Crimes, T.P., and Marcos, A., 1976, Trilobite traces and the age of the lowest part of the Ordovician reference section for NW Spain: Geological Magazine, v. 113, p. 350–356. Davidson, T., 1880, On the species of Brachiopoda that characterize the “Grès Armoricain” of Brittany together with a few observations on Budleigh Salterton “Pebbles”: Geological Magazine, v. 2, p. 337–343. Durand, J., 1985, Le Grés Armoricain. Sédimentologie-Trace Fossils. Milieux de dépôt: Centre Armoricain d’Étude Structurale des Socles: Mémoires et Documents, v. 3, p. 1–150. Emig, C.C., and Gutiérrez-Marco, J.C., 1997, Signification des niveaux à lingulidés à la limite supérieure du Grès Armoricain (Ordovicien, Arenig,
295
Sud-Ouest de l’Europe): Geobios, v. 30, p. 481–495, doi: 10.1016/S00166995(97)80116-7. Fernández-Suárez, J., Gutiérrez-Alonso, G., Jenner, G.A., and Tubrett, M.N., 1999, Crustal sources in Lower Palaeozoic rocks from NW Iberia: Insights from laser ablation U-Pb ages of detrital zircons: Journal of the Geological Society of London, v. 156, p. 1065–1068. Fernández-Suárez, J., Gutiérrez-Alonso, G., Cox, R., and Jenner, G.A., 2002, Assembly of the Armorica microplate: A strike-slip terrane delivery? Evidence from U-Pb ages of detrital zircons: Journal of Geology, v. 110, p. 619–626, doi: 10.1086/341760. Fischer, R.V., and Schmincke, H.U., 1984, Pyroclastic rocks: Berlin, SpringerVerlag, 472 p. Fortey, R.A., and Cocks, L.R.M., 2003, Palaeontological evidence bearing on global Ordovician-Silurian continental reconstructions: Earth Science Reviews, v. 61, p. 245–307, doi: 10.1016/S0012-8252(02)00115-0. Gallastegui, G., Aramburu, C., Barba, P., Fernández, L.P., and Cuesta, A., 1992, El Vulcanismo del Paleozoico Inferior de la Zona Cantábrica (NO de España), in Rábano, I., Gutiérrez-Marco, J.C., and Saavedra, J., eds., El Paleozoico inferior de Ibero-América: Badajoz, Spain, Universidad de Extremadura, p. 435–452. Gallastegui, G., Suárez, O., and Cuesta, A., 2004, Magmatismo, in Vera, J.A., ed., Geología de España: Madrid, Sociedad Geológica de España–Instituto Geológico y Minero de España, p. 47–49. García-Ramos, J.C., Aramburu, C., and Brime, C., 1984, Kaolin tonstein of volcanic ash origin in the Lower Ordovician of the Cantabrian Mountains (NW Spain): Trabajos de Geología: Universidad de Oviedo, v. 14, p. 27–33. Gradstein, F.M., Ogg, J.G., and Smith, A.G., eds., 2004, A geologic time scale 2004: Cambridge, Cambridge University Press, 589 p. Guillier, M.A., 1881, Notes sur les lingules du grès armoricain de la Sarthe: Bulletin de la Société Géologique de France, v. 3, p. 372–377. Gutiérrez-Alonso, G., Fernández-Suárez, J., Collins, A.S., Abad, I., and Nieto, F., 2005, Amazonian Mesoproterozoic basement in the core of the IberoArmorican arc: 40Ar/39Ar detrital mica ages complement the zircon’s tale: Geology, v. 33, p. 637–640, doi: 10.1130/G21485.1. Gutiérrez-Marco, J.C., and Bernárdez, E., 2003, Un tesoro geológico en la Autovía del Cantábrico. El Túnel Ordovícico del Fabar, Ribadesella (Asturias): Madrid, Ministerio de Fomento, 398 p. Gutiérrez-Marco, J.C., and Rodríguez, L., 1987, Descubrimiento de graptolitos arenigienses en la Escama de Rioseco (zona Cantábrica, N de España): Cuadernos do Laboratório Xeolóxico de Laxe, v. 11, p. 209–220. Gutiérrez-Marco, J.C., Babin, C., and Porro Mayo, T., 1997, Moluscos bivalvos de las facies cuarcíticas del Ordovícico Inferior centroibérico: Geogaceta, v. 22, p. 85–88. Gutiérrez-Marco, J.C., Robardet, M., Rábano, I., Sarmiento, G.N., San José Lancha, M.A., Herranz Araújo, P., and Pieren Pidal, A.P., 2002, Ordovician, in Gibbons, W., and Moreno, T., eds., The geology of Spain: London, Geological Society of London, p. 31–49. Gutiérrez-Marco, J.C., Bernárdez, E., Rábano, I., Sarmiento, G.N., Sendino, M.C., Albani, R., and Bagnoli, G., 2003, Ordovician on the move: Geology and paleontology of the “Túnel Ordovícico del Fabar” (Cantabrian free highway A-8, N Spain): Instituto Superior de Correlación Geológica, Serie Correlación Geológica, v. 17, p. 71–77. Hamilton, M.A., and Murphy, J.B., 2004, Tectonic significance of a Llanvirn age for the Dunn Point volcanic rocks, Avalon terrane, Nova Scotia, Canada: Implications for the evolution of the Iapetus and Rheic oceans: Tectonophysics, v. 379, p. 199–209, doi: 10.1016/j.tecto.2003.11.006. Harper, D.A.T., and Owen, A.W., 1984, The Caradoc brachiopod and trilobite fauna of the upper Kirkerud group, Hadeland, Norway: Geologica et Palaeontologica, v. 18, p. 21–51. Heinz, W., Loeschke, J., and Avra, G., 1985, Phreatomagmatic volcanism during the Ordovician of the Cantabrian Mountains (NW Spain): Geologische Rundschau, v. 74, no. 3, p. 623–639, doi: 10.1007/BF01821217. Helbing, H., and Tiepolo, M., 2005, Age determination of Ordovician magmatism in NE Sardinia and its bearing on Variscan basement evolution: Journal of the Geological Society of London, v. 162, p. 689–700, doi: 10.1144/0016-764904-103. Henry, J.-L., 1980, Trilobites ordoviciens du Massif Armoricain: Mémoire de la Société Géologique et Minéralogique de Bretagne, v. 22, p. 1–250. Hinckley, D.N., 1962, Variability in “crystallinity” values among the kaolin deposits of the coastal plain of Georgia and South Carolina: Clays and Clay Minerals, v. 11, p. 229–235, doi: 10.1346/CCMN.1962.0110122.
296
Gutiérrez-Alonso et al.
Huff, W.D., Bergström, S.M., and Kolata, D.R., 1992, Gigantic Ordovician volcanic ash fall in North America and Europe: Biological, tectonomagmatic, and event-stratigraphic significance: Geology, v. 20, p. 875–878, doi: 10.1130/0091-7613(1992)020<0875:GOVAFI>2.3.CO;2. Jaffey, A.H., Flynn, K.F., Glendenin, L.E., Bentley, W.C., and Essling, A.M., 1971, Precision measurements of half-lives and specific activities of 235U and 238U: Physical Review C: Nuclear Physics, v. 4, p. 1889–1906, doi: 10.1103/PhysRevC.4.1889. Kolb, S., and Wolf, R., 1979, Distribution of Cruziana in the Lower Ordovician sequence of Celtiberia (NE Spain) with a revision of the Cruziana rugosagroup: Neues Jahrbuch für Geologie und Paläontologie Monatshefte, v. 1979, p. 457–474. Krogh, T.E., 1982, Improved accuracy of U-Pb ages by the creation of more concordant systems using an air abrasion technique: Geochimica et Cosmochimica Acta, v. 46, p. 637–649, doi: 10.1016/0016-7037(82)90165-X. Kuhnt, W., Hess, S., Holbourn, A., Paulsen, H., and Salomon, B., 2005, The impact of the 1991 Mt. Pinatubo eruption on deep sea communities: A model for the K/T boundary? Palaeogeography, Palaeoclimatology, Palaeoecology, v. 224, p. 83–107, doi: 10.1016/j.palaeo.2005.03.042. Landing, E., Bowring, S.A., Davideck, K.L., Rutshton, A.W.A., Fortey, R.A., and Wimbledon, W.A.P., 2000, Cambrian-Ordovician boundary age and duration of the lowest Ordovician Tremadoc Series based on U-Pb zircon dates from Avalonian Wales: Geological Magazine, v. 137, p. 485–494, doi: 10.1017/S0016756800004507. Linnemann, U., and Romer, R.L., 2002, The Cadomian orogeny in SaxoThuringia, Germany: Geochemical and Nd-Sr-Pb isotopic characterization of marginal basins with constraints to geotectonic setting and provenance: Tectonophysics, v. 352, p. 33–64, doi: 10.1016/S0040-1951(02)00188-9. Loeschke, J., and Zeidler, N., 1982, Early Palaeozoic sills in the Cantabrian Mountains (Spain) and their geotectonic environment: Neues Jahrbuch Geologie und Paläontologie Monatshefte, v. 7, p. 419–439. Ludwig, K.R., 1998, On the treatment of concordant uranium-lead ages: Geochimica et Cosmochimica Acta, v. 62, no. 4, p. 665–676, doi: 10.1016/ S0016-7037(98)00059-3. Ludwig, K.R., 2003, User’s manual for Isoplot 3.00: A geochronological toolkit for Microsoft Excel: Berkeley, California, Berkley Geochronology Center Special Publication 4, 70 p. Murphy, J.B., Gutiérrez-Alonso, G., Nance, R.D., Fernández-Suárez, J., Keppie, J.D., Quesada, C., Strachan, R.A., and Dostal, J., 2006. Origin of the Rheic Ocean: Rifting along a Neoproterozoic suture? Geology, v. 34 p. 325–328. Paris, F., 1981, Les Chitinozoaires dans le Paléozoïque du Sud-Ouest de l’Europe (Cadre géologique-Etude systématique-Biostratigraphie): Mémoires de la Société Géologique et Minéralogique de Bretagne, v. 26, p. 1–496. Paris, F., 1990, The Ordovician chitinozoan biozones of the northern Gondwana domain: Review of Palaeobotany and Palynology, v. 66, p. 181–209, doi: 10.1016/0034-6667(90)90038-K. Paris, F., Robardet, M., Durand, J., and Noblet, C., 1982, The Lower Palaeozoic transgression in southwestern Europe: Palaeontological Contributions of the University of Oslo, v. 280, p. 41. Paris, F., Achab, A., Asselin, E., Chen, X.-h., Grahn, Y., Nolvak, J., Obut, O., Samuelsson, J., Sennikov, N., Vecoli, M., Verniers, J., Wang, X.-f., and Winchester-Seeto, T., 2004, Chitinozoans, in Webby, B.D., Paris, F., Droser, M.L., and Percival, I.G., eds., The great Ordovician biodiversification event: New York, Columbia University Press, p. 294–311. Pickerill, R.K., Romano, M., and Meléndez, B., 1984, Arenig trace fossils from the Salamanca area, western Spain: Geological Journal, v. 19, p. 249–269. Prigmore, J.K., Butler, A.J., and Woodcock, N.H., 1997, Rifting during separation of eastern Avalonia from Gondwana: Evidence from subsidence analysis: Geology, v. 25, p. 203–206, doi: 10.1130/0091-7613(1997)025<0203: RDSOEA>2.3.CO;2. Quesada, C., 1991, Geological constraints on the Paleozoic tectonic evolution of tectonostratigraphic terrenes in the Iberian Massif: Tectonophysics, v. 185, p. 225–245, doi: 10.1016/0040-1951(91)90446-Y. Romano, M., 1991, Lower to Middle Ordovician trace fossils from the Central Iberian zone of Portugal and Spain, in Barnes, C.R., and Williams, S.H., eds., Advances in Ordovician Geology: Ottawa, Ontario, Geological Survey of Canada Paper 90-9, p. 191–204.
Ross, C.A., and Ross, J.R.P., 1995, North American sequences and correlations, in Cooper, J.D., Droser, M.L., and Finney, S.C., eds., Ordovician odyssey: Fullerton, California, Pacific Section, Society of Economic Paleontologists and Mineralogists 77, p. 309–313. Rouault, M., 1850, Note préliminaire sur une nouvelle formation découverte dans le terrain silurien inférieur de la Bretagne: Bulletin de la Société Géologique de France, v. 2, p. 724–744. Sá, A.A., Meireles, C., Coke, C., and Gutiérrez-Marco, J.C., 2003, Reappraisal of the Ordovician stratigraphy and paleontology of Trás-os-Montes (Central-Iberian zone, NE Portugal): Instituto Superior de Correlación Geológica, Serie Correlación Geológica, v. 17, p. 131–136. Sadler, P.M., and Cooper, R.A., 2004. Calibration of the Ordovician timescale, in Webby, B.D., Paris, F., Droser, M.L., and Percival, I.G., eds., The great Ordovician biodiversification event: New York, Columbia University Press, p. 48–51. Salman, K., 2004, The timing of the Cadomian and Variscan cycles in the OssaMorena zone, SW Iberia: Granitic magmatism from subduction to extension: Journal of Iberian Geology, v. 30, p. 119–132. San José, M.A., Rábano, I., Herranz, P., and Gutiérrez-Marco, J.C., 1992, El Paleozoico Inferior de la Zona Centroibérica meridional, in GutiérrezMarco, J.C., Saavedra, J., and Rábano, I., eds., Paleozóico Inferior de IberoAmérica: Badajoz, Spain, Universidad de Extremadura, p. 505–521. Schätz, M., Reichmann, T., Tait, J., Bachtadse, V., Bahlburg, H., and Martin, U., 2002, The Early Palaeozoic break-up of northern Gondwana, new palaeomagnetic and geochronological data from the Saxothuringian basin, Germany: International Journal of Earth Sciences, v. 91, p. 838–849, doi: 10.1007/s00531-002-0261-3. Seilacher, A., 1970, Cruziana stratigraphy of “non-fossiliferous” Palaeozoic sandstones, in Crimes, T.P., and Harper, J.C., eds., Trace fossils: Geological Journal, Special Issue, v. 3, p. 447–476. Seilacher, A., 1992, An updated Cruziana stratigraphy of Gondwanan Palaeozoic sandstones, in Salem, M.J., ed., The geology of Libya, v. 4: Amsterdam, Elsevier, p. 1565–1581. Simancas, J.F., Expósito, I., Azor, A., Martínez Poyatos, D.J., and González Lodeiro, F., 2003, From the Cadomian orogenesis to the Early Palaeozoic Variscan rifting in southwest Iberia: Journal of Iberian Geology, v. 30, p. 53–71. Spjeldnaes, N., 1961, Ordovician climatic zones: Norsk Geologisk Tidskrifft, v. 41, p. 45–77. Stacey, J.S., and Kramers, J.D., 1975, Approximation of terrestrial lead isotope evolution by a two stage model: Earth and Planetary Science Letters, v. 26, p. 207–221, doi: 10.1016/0012-821X(75)90088-6. Suárez, O., Gallastegui, G., Cuesta, A., Corretgé, L.G., and Tarrío, L., 1993, Geoquímica de las rocas volcánicas del Cabo Peñas (Zona Cantábrica): V Congreso de Geoquímica de España. Actas del Congreso, Soria, 21–24 September, p. 42–47. Valverde-Vaquero, P., and Dunning, G.R., 2000, New U-Pb ages for Early Ordovician magmatism in central Spain: Journal of the Geological Society of London, v. 157, p. 15–26. Valverde-Vaquero, P., Marcos, A., Farias, P., and Gallastegui, G., 2005, U-Pb dating of Ordovician felsic volcanism in the Schistose domain of the Galicia-Trás-os-Montes zone near Cabo Ortegal (NW Spain): Geologica Acta, v. 3, p. 27–37. Van der Voo, R., 1988, Palaeozoic paleogeography of North America, Gondwana, and intervening displaced terranes: Comparisons of palaeomagnetism within palaeoclimatology and biogeographical patterns: Geological Society of America Bulletin, v. 100, p. 311–324, doi: 10.1130/00167606(1988)100<0311:PPONAG>2.3.CO;2. Walter, R., 1982, Beiträge zur Kenntnis des Präkambriums und Altpaläozoikums der Montes de Guadalupe/Zentralspanien. I. Stratigraphische und tektonische Übersicht: Münstersche Forschungen zur Geologie und Paläontologie, v. 56, p. 1–28. Webby, B.D., Cooper, R.A., Bergström, S.M., and Paris, F., 2004, Stratigraphic framework and time slices, in Webby, B.D., Paris, F., Droser, M.L., and Percival, I.G., eds., The great Ordovician biodiversification event: New York, Columbia University Press, p. 41–47. MANUSCRIPT ACCEPTED BY THE SOCIETY 3 OCTOBER 2006
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Geological Society of America Special Paper 423 2007
Contrasting mantle sources and processes involved in a periGondwanan terrane: A case study of pre-Variscan mafic intrusives from the autochthon of the Central Iberian Zone Miguel López-Plaza* Mercedes Peinado Francisco-Javier López-Moro M. Dolores Rodríguez-Alonso Asunción Carnicero M. Piedad Franco Juan Carlos Gonzalo Departamento de Geología, Universidad de Salamanca, Plaza de los Caídos s/n, 37008 Salamanca, Spain Marina Navidad Departamento de Petrología, Universidad Complutense, Facultad de Geología, José Antonio Novais 2, 28040 Madrid, Spain
ABSTRACT The innermost domain of the autochthon of the Iberian Massif (Central Iberian Zone) consists of Upper Proterozoic–Lower Cambrian metasedimentary rocks and Early Paleozoic augen gneisses. The former were intruded by mafic magmas as small bodies that later became amphibolites under Variscan metamorphism, gabbro-diabasic textures being sometimes well preserved. Three main groups of amphibolites can be established: (A) the light rare earth element (LREE)-depleted group, characterized by an extremely low rare earth element (REE) fractionation factor [(La/Lu)CN = 0.27–0.34] and low Ti content; (B) the flat REE pattern group, characterized by a small REE fractionation factor [(La/Lu)CN = 0.95–1.25]; and (C) the LREE-enriched group, characterized by a strong fractionation factor [(La/Lu)CN = 3.53–15.04] and high Ti content. Regular major element variations for the depleted amphibolites point to a low-pressure fractional crystallization as the major process, the strongly depleted amphibolites considered to be parent magmas, although crustal contamination accounts for their broad Nd and Sr isotope ranges as well as trace element variations. Part of the enriched amphibolite samples have an ocean island basalt–like signature and plot within the mantle array, suggesting a mantle mixed source. Other samples show decoupling of Nd and Sr isotopic systems as a response to a probable mixing process involving fluid-rock interaction during subsolidus evolution, which is supported by their high δ18O values. Geochemical characterization of amphibolite groups *E-mail:
[email protected]. López-Plaza, M., Peinado, M., López-Moro, F.-J., Rodríguez-Alonso, M.D., Carnicero, A., Franco, M.P., Gonzalo, J.C., and Navidad, M., 2007, 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, 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. 297–313, doi: 10.1130/2007.2423(14). For permission to copy, contact
[email protected]. ©2007 Geological Society of America. All rights reserved.
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López-Plaza et al. is consistent with an extensional within-plate tectonic setting, whereas processes involved may account for a similar mantle-crust interaction that occurred throughout the entire Central Iberian Zone and probably all over the northern margin of Gondwana during late Neoproterozoic–Early Paleozoic times. Keywords: Amphibolites, Early Paleozoic, extensional event, Central Iberian Zone, mantle sources, AFC process, mantle-crust interplay
INTRODUCTION The Iberian Massif includes a variety of units of different paleogeographic origin (Quesada, 1991), the Cantabrian, West Asturian-Leonese, and Central Iberian Zones representing continental fragments of the Iberian autochthon (Neoproterozoic Gondwana) (Fig. 1). Differences between them in their tectonothermal evolution can be related to their respective positions during Variscan evolution. The Late Proterozoic–Early Cambrian magmatism associated with the Cadomian orogeny is relatively scarce in the Iberian autochthon (Fernández-Suárez et al., 1998; Rodríguez-Alonso et al., 2004b), although Cadomian magmatism is a widespread feature in the so-called “Ossa-Morena terrane” in southwest Iberia, which is considered to be a continental arc accreted during Upper Proterozoic–Lower Paleozoic times to the Iberian autochthon (Quesada et al., 1991; Bandrés et al., 2002; Murphy et al., 2002; Sánchez-García et al., 2003, and references therein). The most internal parts of the Iberian Massif (Central Iberian Zone, Spain) consist of Upper Proterozoic–Lower Cambrian metasedimentary rocks and Early Paleozoic augen gneisses. The former were intruded by small bodies of mafic magma that later became amphibolites. There is general consensus about the meaning of the Early Paleozoic magmatism throughout the Variscides of Europe in terms of rift-related bimodality in a pre-Variscan extensional scenario (Crowley et al., 2000). Amphibolites representing the mafic member have been reported for the Iberian Massif (Escuder and Navidad, 1999) and are thought to have been derived from several sources as a result of evolving rift zones with a crust of varying thickness (Crowley et al., 2000). Pervasive Late Proterozoic and Early Paleozoic mafic magmatism in the accreted terrane of Ossa-Morena could have resulted from an extreme crustal thinning during the Cambrian-Ordovician rifting event (SánchezGarcía et al., 2003). Owing to a lack of other petrotectonic indicators, many authors consider that geochemical discrimination criteria applied to amphibolites may be extremely valuable in reconstructing the tectonic setting, despite a strong Variscan overprinting and superimposed petrogenetic processes. The difficulties are more evident in high-grade metamorphic areas, such as those considered here, in which a lack of knowledge about any of the sequential intrusive events renders the issue even more problematic. Work on amphibolites from the Central Iberian Zone (Beetsma, 1995; Escuder and Navidad, 1999; López-Plaza and
López-Moro, 1999; Barbero and Villaseca, 2000; Ferreira et al., 2000; López-Moro, 2000; Peinado et al., 2002) has focused on geochemical characterization and discrimination. Thus, the main goal of the present study emphasizes the role played by petrogenetic processes in order to shed light on the meaning of the geochemical variations, hence allowing magmatic and overprinted processes to be discriminated, as well as possible artifacts. This work provides an opportunity for the Early Paleozoic mafic magmatism from the Central Iberian and the Ossa-Morena Zones to be compared in the light of mantle-crust interplay and also addresses the extent to which homogeneity existed in the extensional continental margin of Gondwana. Both tholeiitic and alkaline affinities have been reported for amphibolites from the Central Iberian Zone and, mainly, the Ossa-Morena Zone (Gómez-Pugnaire et al., 2003; SánchezGarcía et al., 2003; Pereira et al., 2004), where different sources appear to have been involved, including the ocean island basalt (OIB)-affinity asthenospheric mantle (Gómez-Pugnaire et al., 2003; Sánchez-García et al., 2003). Estimates of the ages of the protoliths of the amphibolites from the Ossa-Morena Zone were obtained using the U-Pb method on zircons (Ordóñez-Casado, 1998), permitting two main groups to be distinguished: (1) Upper Precambrian amphibolites, ranging from 596 to 577 Ma; and (2) Lower Paleozoic amphibolites, with ages from 525 to 483 Ma. Such a wide range hinders establishment of any precise correlation with the amphibolites studied in this work. Other crustal blocks of the northern margin of Gondwana contain amphibolitic rocks that appear to be good indicators of the tectonic setting during late Neoproterozoic–Early Paleozoic times. Alkaline metabasalts from the Bohemian Massif (Floyd et al., 2000) show a geochemical similarity to that of the amphibolite group established in this work, allowing us a correlation not only in terms of chemical characterization and magmatic processes (incompatible element enrichment), but also in relation to tectonic setting (intracratonic rifting). GEOLOGICAL SETTING The Central Iberian zone represents the innermost part of the Iberian Variscan belt in a similar way to the zones of the South Armorican Zone and the Central Massif in France. A subzone made up of allochthonous and suballochthonous terranes has been distinguished in the northwest of the Central Iberian Zone: the Galicia Media-Tras-os Montes subzone (Farias et al., 1987; Fig. 1). The autochthonous terrane, like comparable terranes in
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Figure 1. Geological map of the autochthonous terrane of Central Iberian Zone, showing locations of amphibolite samples on the Orthogneiss domain. Inset: reconstruction of Pangea in the area surrounding Iberia, based on Lefort (1989), Skehan and Rast (1995), and Martínez Catalán et al. (1999).
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southern Europe, including the Ossa-Morena Zone, was part of the active northern margin of Gondwana during Upper Proterozoic–Lower Paleozoic times (Murphy et al., 2000; GutiérrezAlonso et al., 2003; Rodríguez-Alonso et al., 2004b). Within the autochthonous terrane of the Central Iberian Zone, an Orthogneiss domain can be distinguished in which several types of augen gneiss outcrop, either as a 600-km-long belt (Ollo de Sapo antiform) or as mega-enclaves within anatectic Variscan granites. The Orthogneiss domain has been mapped taking into account both the occurrences of plutonometamorphic belts and augen gneiss outcrops (Fig. 1). As shown in Figure 1, amphibolites outcrop throughout the Orthogneiss domain. They are in a Neoproterozoic–lowermost Cambrian metasedimentary succession that took place in the northern margin of Gondwana between the end of the Cadomian cycle and the beginning of the Cambro-Ordovician cycle, recording the passage from an active margin regime during the Neoproterozoic-Cambrian limit to the rifting and fragmentation of this margin, which became passive in Cambrian-Ordovician times (Rodríguez-Alonso et al., 2004a,b), as occurred in other marginal areas of northern Gondwana (Murphy and Nance, 1989; Nance and Murphy, 1994). These metasedimentary rocks have always been considered to form a Schist-Graywacke Complex in which two main informal units can be described: the Lower Unit and the Upper Unit, interpreted as a continuous process of siliciclastic sedimentation with the development of turbiditic facies, which evolved upward into mixed siliciclastic-carbonate slope-platform deposits within a tectonic-controlled basin and associated with volcanic activity (Rodríguez-Alonso and Alonso Gavilán, 1995, and references therein; Rodríguez-Alonso et al., 2004b). Amphibolites appear in the Upper Unit sequence, mainly constituted by metapelites with minor intercalations of marbles, calc-silicate rocks, and tonalitic gneisses, the whole sequence having been affected by high-grade metamorphic conditions (Bellido et al., 1981; López-Moro and López-Plaza, 1992; Gonzalo et al., 1994; Barbero and Villaseca, 2000). In the Tormes Dome area, a sequence from bottom to top has been defined as follows: augen gneisses, fine-grained gneisses, and the complex group (Schist-Graywacke Complex) mentioned above. Within this ensemble, the amphibolites outcrop in association with tonalitic gneisses and metapelites, the latter of which become more abundant toward the top of the sequence (Martínez, 1974; LópezMoro and López-Plaza, 1992). In the Central System, a similar sequence has been reported, with dominant metapelites and discontinuous layers of marbles and calc-silicate rocks (Capote et al., 1977; Barbero and Villaseca, 2000, and references therein). It is inferred from these descriptions that a Neoproterozoic–lowermost Cambrian metasedimentary sequence was intruded by felsic magmas (augen gneisses) during Cambrian-Ordovician extensional events. Geochronological data obtained from orthogneissic rocks point to a relatively wide time span, from late Neoproterozoic-Cambrian times (546 Ma, U-Pb zircon, after Zeck et al., 2004; 526 Ma, U-Pb zircon, after Ferreira et al., 2000) to Ordovician times (487 Ma, U-Pb zircon, after Valverde Vaquero
and Dunning, 2000). As far as we know, neither amphibolites nor other mafic rocks have been found within the orthogneisses. Because no geochronological data for the amphibolites are currently available, in this work a minimum age of 546 Ma is considered as a reference for these rocks. The amphibolites occur as meter- to decameter-scaled dikelike bodies or, more commonly, as sheet-like discontinuous intrusions into the metasedimentary sequence. They have a relatively massive appearance, modified by metamorphic foliation, although local chilled margins have been recognized, and gabbro-diabasic textures have been sometimes preserved. In this work three main groups of amphibolites are established: (A) the light rare earth element (LREE)-depleted group, (B) the flat rare earth element (REE) pattern group, and (C) the LREE-enriched group. No significant differences in field observations have been found for the three groups, and the latter two may appear within the same area (e.g., the Tormes Dome area; Fig. 1). It is worth noting that a large proportion of the samples are located in the Tormes Dome area: four out of seven samples for the depleted amphibolites, and eleven out of fourteen samples for the enriched amphibolites. This circumstance provides a twofold reliability for the protoliths: on one hand in terms of magmatic processes within the Tormes Dome itself, and on the other hand for the entire Orthogneissic domain, because the rest of the samples display either a geochemical affinity or a complementary nature, which is the case for the strongly depleted amphibolites. Taken together, these properties define distinct, coherent groups, as will be shown in this work. ANALYTICAL METHODS Eighteen fresh samples were selected for the analysis of whole-rock; five of them for electron microprobe. Minerals were analyzed with a Camebax (CAMECA) microprobe at the University of Oviedo, Spain. Operating conditions were 15 kV and 20 nA, with a counting time of 20 s. A few apatite-bearing samples were selected for analysis of cathodoluminescence at the Laboratory of the University of Salamanca. Whole-rock analyses were carried out at the ACTLABS in Ancaster, Ontario, Canada. Major element analyses were carried out by inductively coupled plasma-atomic emission spectroscopy (ICP-AES), whereas trace elements and REE were determined by ICP-mass spectrometry (ICP-MS). Nine samples were also selected for the study of Rb-Sr and Sm-Nd isotopes. They were analyzed at the Center of Isotopic Geochemistry and Geochronology of the Universidad Complutense de Madrid. Sr was run on Ta single filaments, whereas Sm and Nd were run on Ta-Re-Ta triple filaments. All of them were determined using thermal ionization on a VG Sector 54® mass spectrometer, with five Faraday cups, by means of a multicollector thermal ionization mass spectrometer with data acquired in multidynamic mode, except for Sm, which was measured in simple collection. The measurements of Sr were corrected for possible interference by 87Rb and normalized to
Contrasting mantle sources and processes involved in a peri-Gondwanan terrane 88
Sr/86Sr = 0.1194. The measurements of Nd were corrected for interference by 142Ce and 144Sm and normalized to 146Nd/144Nd = 0.7219. During this study, the NBS-987 standard gave an average value (n = 12) for 87Sr/86Sr of 0.710271 ± 0.00002, with an average value (n = 7) for 143Nd/144Nd of 0.511806 ± 0.000005, which is equivalent to those obtained by this laboratory in a period of 24 months for 87Sr/86Sr = 0.710261 ± 0.00002 (2σ, n = 112) and 143 Nd/144Nd = 0.511809 ± 0.000008 (2σ, n = 41). The εNd values were calculated using the following bulk Earth parameters: 143 Nd/144Nd = 0.512638; 147Sm/144Nd = 0.1967. The 2σ standard deviation error in εNd calculations is ± 0.4. Four oxygen isotope analyses were performed at the University of Salamanca. Oxygen extraction for isotope analysis used the technique of Clayton and Mayeda (1963) but employed a loading technique similar to that described by Friedman and Gleason (1973) and used CIF3 as reagent (Borthwick and Harmon, 1982). About 10 mg of finely powdered sample was loaded into nickel vessels and reacted for 15 h at 690 °C. The oxygen released was converted to CO2 by means of a carbon rod heated by a platinum wire. Isotope ratios were determined on a VG SIRA-II® mass spectrometer. Results are reported in the usual notation as δ18O values relative to the V-Standard Mean Ocean Water reference standard. The typical overall reproducibility of duplicate runs was within ± 0.2‰ (1σ). PETROGRAPHY The rocks have been strongly affected by Variscan metamorphism; hence, no primary minerals can be found. The samples studied show granonematoblastic microstructures, sometimes porphyroclastic textures, and, more rarely, subhedral plagioclase crystals. Their average mineral compositions are: hornblende [(Na,K)0.637Ca1.652(Mg,Fe,Ti)4.545Al2.387Si6.225]; plagioclase (An 35–54); ±diopside (wollastonite47.6; enstatite30.3; ferrosilite22.01); ±garnet (almandine54–58.17; grossular18.61–17.94; pyrope11.19–13.81; spessartine9.8–6.06); and ±biotite, ilmenite, titanite, apatite, zircon, and scarce rutile. Although there are some differences among the types, the strongly depleted rocks maintain a gabbro-diabasic microstructure at the border of the tabular-shaped intrusions, displayed by subhedral plagioclase, which is prismatic, although internally transformed into a granoblastic aggregate. These rocks contain metamorphic garnet surrounded by granoblastic plagioclase with small green spinel, the original ferromagnesian minerals being transformed into green amphibole. The depleted group shows the simplest metamorphic mineralogy, mainly consisting of amphibole and plagioclase and showing an overall fine-grained, granonematoblastic texture. Besides plagioclase and amphibole, the enriched group shows metamorphic clinopyroxene and quite often biotite, although only in the more potassic compositions. Titanite is also abundant. Most of the mineral assemblages reflect high- and medium-temperature conditions and late facies of Variscan metamorphism; although
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some high-pressure first-stage metamorphic relics have been preserved (Barbero and Villaseca, 2000) in the depleted group. GEOCHEMICAL GROUPS The major element contents indicate that most of the rocks analyzed are basalts in composition with several basaltic andesites (Tables 1 and 2), although more evolved rocks, such as dacites, have been reported by Barbero and Villaseca (2000). Their mgN ranges between 0.702 and 0.350, suggesting that these magmas evolved from more primitive types, as can be deduced by bivariate correlations, such as decreasing Cr and CaO/Al2O3 ratios, with differentiation. Trace and REE contents allow three different groups of mafic rocks to be established (Figs. 2 and 3; Tables 1–3): A. Strongly depleted amphibolites, characterized by an extremely low REE content (∑REE = 20.4–21.3). These have positive trends consistent with fractionation of the LREE relative to the heavy ones [(La/Lu)CN = 0.28–0.34]. Also, they have very high Zr/Nb (62–56) and La/Nb (2.6–2.2) ratios, a low Ti content (0.58–0.66 TiO2%), and high Cr and Ni levels. The multi-element patterns show a strong negative Nb anomaly; a moderate negative anomaly in P, Zr, and Hf; and a strongly positive Sr anomaly. At present, only one outcrop of this kind has been recorded; a decameter-sized body intruded into migmatized metasediments in the southern slope of the Spanish Central system (Fig. 1). B. Depleted amphibolites, characterized by a flat REE pattern [(La/L u)CN = 0.95–1.25; ΣREE = 67–92]; a slight Eu anomaly (0.81–1.07); an intermediate Ti content (1.72– 2.14 TiO2%); a slight negative anomaly in Nb, Ta, Sr; and intermediate Zr/Nb (37.1–22.0) and La/Nb (2.09–1.22) ratios. They also display a flat multi-element pattern, with moderate Nb and Ta troughs. The rocks studied here (four samples) are localized in the Tormes Dome region (Fig. 1). Similar flat-patterned amphibolites have been described in other localities in the Central Iberian Zone, such as in the Spanish Central System around Segovia (Barbero and Villaseca, 2000), some of which have been included in this work (two samples), as well as one sample from the Caramulo area, Portugal (Beetsma, 1995; Fig. 1), which was considered of Lower Paleozoic age. Other similar amphibolites have been reported in Portugal, Foz do Douro, close to the Ossa-Morena Zone, which, based on their model ages, were interpreted as representative of a mid-ocean ridge basalt (MORB)-type magmatism related to a prior oceanization period (Leterrier and Noronha, 1998). In any case, these patterns suggest that this group is comparable to the late tholeiitic amphibolites from the whole of the Ossa-Morena Zone (Sánchez-García et al., 2003), and particularly with respect to those from the Evora Massif in Portugal (Pereira et al., 2004) and those from Badajoz-Córdoba (Group 3, after Gómez-Pugnaire et al., 2003) dated at between 483 and 525 Ma.
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Sample
TABLE 1. MAJOR AND TRACE ELEMENTS OF REPRESENTATIVE DEPLETED AND STRONGLY DEPLETED AMPHIBOLITES BR-15 BR-16 DT 134 DT 139 DT 135 DT 139B 96895* 96896*
(wt%) SiO2 47.49 48.55 48.10 49.01 51.14 52.26 46.34 48.23 TiO2 0.66 0.58 1.72 2.14 1.95 2.08 3.81 3.46 Al2O3 17.52 15.78 15.85 15.16 15.13 14.8 12.82 13.12 Fe2O3t 9.21 9.22 10.60 12.43 11.33 15.04 18.66 17.75 MnO 0.15 0.17 0.24 0.26 0.86 0.23 0.31 0.28 MgO 9.66 10.12 7.64 6.98 6.66 5.30 4.77 4.19 CaO 12.58 10.98 11.09 10.42 7.49 9.07 9.54 8.95 Na2O 1.77 2.62 2.43 2.31 3.57 0.70 3.33 3.39 K2O 0.05 0.16 0.34 0.38 0.21 0.23 0.15 0.36 P2O5 0.03 0.05 0.16 0.26 0.15 0.26 0.45 0.52 LOI 0.92 1.12 1.14 0.96 0.80 0.18 n.d. n.d. Total 100.04 99.35 99.31 100.31 99.29 100.15 100.18 100.25 (ppm) Cs 1.8 1.4 2.0 b.d.l. 3.0 b.d.l. n.d. n.d. Rb 3.0 4.0 4.7 6.0 11.0 4.0 4.0 15.0 Sr 142.0 269.0 210.0 189.0 78.0 52.0 62.0 76.0 Ba 7.0 71.0 152.0 244.0 38.0 18.0 25.0 67.0 Cr 439.0 458.0 272.0 222.0 251.0 126.0 n.d. n.d. Ni 145.0 158.0 79.0 71.0 92.0 44.0 n.d. n.d. Sc 36.0 37.0 41.8 42.0 39.2 47.0 n.d. n.d. Co 49.0 42.0 47.9 48.0 53.7 59.0 n.d. n.d. V 209.0 207.0 283.0 335.0 307.0 358.0 n.d. n.d. Cu 51.0 38.0 72.0 35.0 99.0 27.0 n.d. n.d. Zn b.d.l. 34.0 120.0 107.0 58.0 109.0 n.d. n.d. Y 19.0 17.0 34.0 45.0 44.0 45.0 50.0 63.0 Zr 31.0 28.0 115.0 159.0 139.0 132.0 191.0 241.0 Hf 0.9 0.8 3.3 4.2 3.9 3.8 n.d. n.d. Nb 0.5 0.5 3.1 6.0 6.3 5.0 12.0 15.0 Ta 0.1 b.d.l. 0.27 0.5 0.36 0.5 n.d. n.d. Th 0.2 0.2 0.55 1.1 1.26 1.4 1.3 1.84 U 0.01 b.d.l. 0.78 0.5 1.19 0.5 n.d. n.d. La 1.1 1.3 6.48 10.1 7.72 10.3 9.06 13.04 Ce 2.9 3.0 15.2 23.1 20.8 23.5 23.01 36.66 Pr 0.52 0.49 2.712 3.84 3.157 3.39 n.d. n.d. Nd 3.1 3.1 13.8 18.7 15.6 15.9 16.33 24.61 Sm 1.5 1.4 4.44 5.5 5.15 4.6 5.25 7.9 Eu 0.69 0.64 1.631 1.94 1.427 1.74 1.8 2.52 Gd 2.2 2.1 4.64 7.1 5.43 6.6 7.52 9.59 Tb 0.5 0.4 1.01 1.3 1.18 1.3 n.d. n.d. Dy 3.3 2.9 6.47 7.8 7.28 7.8 8.32 10.96 Ho 0.7 0.6 1.35 1.6 1.56 1.7 n.d. n.d. Er 2.1 2.0 3.96 5.0 4.68 5.2 4.8 6.14 Tm 0.32 0.3 0.634 0.75 0.723 0.8 n.d. n.d. Yb 2.0 1.9 3.75 4.7 4.33 5.0 4.51 5.88 Lu 0.34 0.31 0.55 0.67 0.669 0.71 0.74 0.93 Note: b.d.l.—below detection limit; Fe2O3t—total iron as ferric; LOI—loss on ignition; n.d.—not determined. *After Barbero and Villaseca (2000). † After Beetsma (1995).
C. Enriched amphibolites, characterized by a LREE-enriched pattern. These have a strong fractionation factor [(La/Lu)CN = 3.53–15.04]; a wide differentiation-controlled Eu anomaly (1.40–0.93); high content in Ti (2.14–3.87 TiO2%), Nb, and Ta; and low Zr/Nb (8.27–4.0) and La/Nb (1.53–0.71) ratios. The multi-element patterns show the highest element content, displaying an overall negative slope. The negative K anomaly relative to Nb and Th is emphasized, as it has been reported in many worldwide volcanic
P3B
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49.10 1.58 15.10 11.00 0.18 5.85 15.09 1.03 0.51 0.16 n.d. 99.6 1.6 13.0 375.0 52.0 234.0 68.0 41.7 51.5 140.0 n.d. 96.0 38.4 121.0 2.85 5.6 n.d. 0.4 0.2 4.7 14.4 n.d. 11.3 3.67 1.49 n.d. 0.85 n.d. n.d. n.d. n.d. 3.26 0.52
provinces (see Späth et al., 2001), suggesting some K-rich mineral in the source. The enriched amphibolites of the Central Iberian Zone have been reported not only in the Central System (Escuder and Navidad, 1999; Peinado et al., 2002) but also in the Tormes Dome region (LópezPlaza and López-Moro, 1999; López-Moro, 2000), as well as in central-northern Portugal (Beetsma, 1995). According to Winchester and Floyd’s diagrams (1977) a subalkaline character for the A and B groups as well as alkaline
Contrasting mantle sources and processes involved in a peri-Gondwanan terrane TABLE 2. MAJOR AND TRACE ELEMENTS OF REPRESENTATIVE ENRICHED AMPHIBOLITES Sample 9929* DT 132 DT 142 DT 133 DT 140 DT 143 DT 141 DT 136 DT 138 DT 142D DT 115 CAS5 (wt%) SiO2 52.45 45.97 44.01 45.81 46.42 46.89 46.94 47.14 47.18 47.59 53.87 51.52 TiO2 2.04 3.05 2.97 2.70 3.87 2.60 3.03 2.27 3.84 3.51 2.14 2.67 Al2O3 15.21 12.44 14.94 16.23 14.94 8.58 15.52 20.38 16.21 13.68 15.66 14.64 Fe2O3t 11.80 12.35 14.15 12.78 13.92 13.93 10.00 9.16 14.10 14.03 11.49 12.77 MnO 0.159 0.20 0.33 0.23 0.30 0.18 0.17 0.12 0.17 0.23 0.16 0.21 MgO 7.00 10.96 6.53 6.09 3.78 15.02 4.48 3.45 5.26 6.70 5.18 7.30 CaO 9.26 10.8 15.39 13.41 15.47 11.23 15.72 12.96 9.66 12.93 9.69 8.33 Na2O 1.70 0.92 0.51 0.59 0.76 0.76 2.47 3.07 1.06 0.91 2.27 1.58 K2O 0.31 0.39 0.48 0.26 0.04 0.26 0.61 0.79 1.30 0.33 0.25 0.28 P2O5 0.26 0.59 0.76 0.38 0.57 0.17 0.46 0.29 0.65 0.49 0.24 0.41 LOI 0.30 1.00 0.77 0.53 0.16 1.24 0.99 1.01 0.54 0.33 0.00 0.80 Total 100.49 98.67 100.84 99.01 100.23 100.86 100.39 100.64 99.97 100.73 100.95 100.5 (ppm) Cs 1.9 3.1 2.5 4.2 13.4 3.7 6.4 7.5 31.1 2.8 0.7 2.1 Rb 9.0 16.0 24.0 5.9 3.0 9.0 18.0 16.0 58.0 4.0 5.0 12.0 Sr 457.0 889.0 1177. 412.0 319.0 329.0 322.0 588.0 860.0 1076.0 324.0 248.0 Ba 152.0 109.0 247.0 93.0 92.0 18.0 117.0 178.0 316.0 261.0 251.0 108.0 Cr 295.0 430.0 224.0 244.0 213.0 690.0 125.0 49.0 28.0 224.0 66.0 198.0 Ni n.d. 254.0 123.0 120.0 132.0 281.0 40.0 28.0 62.0 149.0 53.0 137.0 Sc n.d. 27.8 20.0 33.9 26.0 34.0 33.0 22.0 19.0 25.0 25.0 22.0 Co 52.0 68.2 54.0 75.9 72.0 78.0 58.0 49.0 54.0 64.0 65.0 69.0 V 242.0 269.0 223.0 300.0 348.0 261.0 291.0 236.0 214.0 308.0 220.0 215.0 Cu 117.0 3.0 62.0 37.0 58.0 b.d.l. 68.0 14.0 42.0 b.d.l. 42.0 b.d.l. Zn 92.0 107.0 186.0 105.0 122.0 114.0 81.0 56.0 109.0 147.0 91.0 173.0 Y 25.0 24.0 41.0 23.0 30.0 20.0 31.0 21.0 40.0 36.0 27.0 32.0 Zr 135.0 230.0 425.0 164.0 228.0 103.0 232.0 155.0 320.0 259.0 128.0 176.0 Hf 3.6 5.6 9.3 4.2 5.8 3.6 5.6 3.9 7.6 8.0 3.5 6.2 Nb 21.0 57.0 106.0 31.0 37.0 12.0 28.0 24.0 45.0 36.0 20.0 31.0 Ta 1.7 3.61 7.5 2.16 3.0 1.1 2.4 2.1 3.5 2.8 1.7 1.8 Th 3.7 4.92 10.4 2.3 4.1 1.1 2.2 1.8 3.2 3.6 2.8 2.5 U 1.0 1.42 3.1 0.64 1.2 0.3 1.3 0.6 1.0 1.1 0.7 0.8 La 21.6 43.6 79.6 22.9 32.0 8.9 23.1 17.9 34.2 26.5 17.5 28.4 Ce 42.2 88.7 147. 47.8 66.5 20.1 50.5 38.4 71.7 57.1 35.0 58.9 Pr 4.93 10.6 16.6 6.04 8.48 2.95 6.66 4.94 9.24 7.89 4.43 6.77 Nd 21.7 39.6 63.2 25.1 36.3 14.0 29.0 20.9 40.6 34.5 19.2 34.1 Sm 4.9 8.08 11.5 6.1 8.0 4.0 6.7 4.7 9.5 8.4 4.9 7.2 Eu 1.72 2.52 4.02 1.799 3.1 1.93 2.45 1.82 3.53 3.04 1.9 2.53 Gd 5.0 7.02 10.6 5.43 8.1 4.3 7.0 4.9 9.9 8.3 5.7 7.3 Tb 0.8 0.99 1.5 0.91 1.2 0.7 1.1 0.8 1.5 1.3 1.0 1.3 Dy 4.5 5.33 8.1 5.1 6.2 4.1 6.2 4.2 7.9 7.3 5.3 6.4 Ho 0.9 0.95 1.4 0.94 1.1 0.7 1.1 0.8 1.4 1.3 1.0 1.2 Er 2.3 2.62 4.1 2.54 3.0 2.0 3.3 2.2 3.9 3.5 2.8 3.2 Tm 0.32 0.321 0.54 0.344 0.38 0.24 0.44 0.28 0.49 0.45 0.35 0.39 Yb 2.0 1.84 3.2 1.95 2.3 1.5 2.7 1.8 2.9 2.8 2.3 2.4 Lu 0.28 0.272 0.44 0.273 0.31 0.21 0.37 0.25 0.38 0.36 0.3 0.33 Note: b.d.l.—below detection limit; Fe2O3t—total iron as ferric; LOI—loss on ignition; n.d.—not determined. *After Barbero and Villaseca (2000). † After Beetsma (1995).
affinity for the C group can be inferred from their Nb/Y ratios (Group A: 0.26–0.29; Group B: 0.09–0.14; Group C: 0.60– 2.59; Table 3). Ternary discrimination diagrams support the former types, showing the following apparent magma affinities: normal MORB (N-MORB) for Group A, enriched MORB (E-MORB) for Group B, and within-plate affinity for Group C (Fig. 4). OIB-like affinity for the enriched amphibolites is shown by incompatible-element ratios with similar values to those for OIB (Table 3), in contrast to the depleted and strongly depleted amphibolites.
303
†
SF5
P2F
50.05 2.76 15.95 12.08 0.22 4.36 7.29 3.42 1.38 0.80 1.37 99.68
50.80 3.11 15.3 13.98 0.18 5.26 8.3 2.67 0.12 0.51 n.d. 100.23
4.3 68.0 874.0 242.0 0.0 0.0 13.0 11.0 133.0 b.d.l. 89.0 38.0 280.0 7.4 31.0 4.3 4.8 1.3 47.9 99.0 11.8 51.7 11.0 4.0 10.7 1.5 8.1 1.3 3.6 0.47 2.8 0.39
b.d.l. 3.0 237.0 55.0 58.0 48.0 24.0 57.8 140.0 n.d. 150.0 44.5 335.0 7.79 35.1 2.18 3.8 b.d.l. 26.9 68.2 n.d. 37.7 8.95 3.13 n.d. 1.21 n.d. n.d. n.d. n.d. 2.89 0.43
A range of supposed tectonic environments from spreading center (N-MORB), plume-dominated (P-MORB), and withinplate (alkali-basalts) can be inferred from the geochemical diagrams, in agreement with results on metabasalts from the north Bohemian Massif and Central Massif (Crowley et al., 2000). A major rifting event could be consistent with the OIB-affinity amphibolites (Group C), whereas an intracontinental back-arc setting cannot be completely discarded for the depleted ones (Groups A and B). The latter display similar geochemical characteristics to other metabasalts from the Variscides of Europe
304
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Enriched amphibolites
Enriched amphibolites
Depleted
Depleted
Strongly depleted
Strongly depleted
100
Rock / Chondrite
Rock / Chondrite (except Rb, K, P)
1000
10
100
10
1
1
0,1
La Ce Pr Nd Pm Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
Ba Rb Th K Nb Ta La Ce Sr Nd P Sm Zr Hf Ti Tb Y Tm Yb
Figure 2. Chondrite-normalized multi-element spidergram, after Thompson et al. (1984), renormalized to Yb = 10.
Figure 3. Chondrite-normalized REE patterns. Normalization values after Anders and Grevesse (1989).
TABLE 3. SOME ELEMENTAL RATIOS FOR MANTLE, CRUST, AND AVERAGE VALUES OF AMPHIBOLITES OF THIS WORK U n it
Th/Ta
Th/Nb
Th/La
T h/ U
Ce/Th
Ce/Nb
Zr/Nb
La/Nb
Nb/U
Primitive mantle
2.07
0.12
0.12
4.05
20.88
2.49
15.71
0.96
33.95
Nb/Y 0.16
N-MORB
0.91
0.05
0.05
2.55
62.50
3.22
31.76
1.08
49.57
0.08
OIB
1.48
0.08
0.11
3.92
20.00
1.67
5.83
0.77
47.06
1.66
Bulk continental crust
3.5
0.32
0.22
3.85
9.43
3.00
9.09
1.45
12.09
0.55
Enriched amphibolites (average)
1.26
0.09
0.11
3.26
17.87
1.76
6.51
0.84
33.12
1.21
Strongly depleted amphibolites (average)
2.00
0.40
0.18
20.00
14.50
5.80
59.00
2.40
50.00
0.03
Depleted amphibolites (average)
2.57
0.15
0.12
1.56
21.38
3.21
22.76
1.25
9.33
0.16
Sources: Primitive mantle, normal mid-oceanic-ridge basalt (N-MORB), and oceanic-island basalt (OIB) values after Sun and McDonough (1989); bulk continental crust after Taylor and McLennan (1985).
that have occasionally been interpreted as an active volcanic arc; however, crustal contamination of mantle-derived basaltic rocks may produce an arc-type signature (see discussion in Crowley et al., 2000), as will be argued in the next sections. GEOCHEMICAL DISCRIMINATION TRENDS The ratio of highly incompatible elements whose bulk partition coefficients are very similar provides indications of contrasted igneous processes and probably mantle sources (Table 3). Binary diagrams, such as Th versus U (Fig. 5), reveal different trends for both types of amphibolites, the enriched amphibolites showing a good correlation, in contrast with a rather scattered plot for the depleted ones. Some incompatible element ratio versus incompatible element diagrams, such as the Th/Ta versus Zr diagram (Fig. 6),
permit two contrasting trends to be distinguished, similarly to the metabasites from Bohemian Massif (Floyd et al., 1996; Floyd et al., 2000): 1. The crustal contamination affinity trend is represented by the depleted amphibolites (Group B). This trend is characterized by a dramatic increase in the Th/Ta ratio, combined with a low Th/Tb ratio and a low Zr content. High-level crustal contamination does not appear to be reconciled with its very high enrichment in Zr and Th. This crustal trend is equivalent to the Bohemian low-Ti metatholeiites and is in agreement with middle or lower crust values. 2. The OIB-type source affinity trend is represented by Group C and is similar to the so-called “plume-influence” trend defined by the Bohemian enriched metabasalts. The average OIB composition plots within this trend either
Contrasting mantle sources and processes involved in a peri-Gondwanan terrane
305
2Nb Strongly depleted amphibolite Depleted amphibolites
100
Enriched amphibolites
10
Th (ppm)
WPA
OIB 1
0.1 PM
P-MORB 0.01
WPT
VAB
NMORB
0.001 0.001
N-MORB
0.01
0.1
Y
ΔNb = 1.74 + log(Nb/Y) – 1.92 log(Zr/Y). Accordingly, positive values of ΔNb (from 0.0 to +0.3) in the enriched amphibolites (Group C) appear to be a good indicator of an Icelandic-like mantle source. In contrast, negative values of ΔNb, from 0.0 to –0.3, in the depleted amphibolites (Group B) strongly suggest an upper mantle source, accounting for the low Nb/Y trend. Contamination of N-MORB–like (Nb/Y = 0.08; see Table 3) magma with continental crust (Nb/Y = 0.55) can never produce magmas with positive values of ΔNb, and the enriched amphibolites (Nb/Y = 1.21) essentially reflect an OIB-like source (Nb/Y = 1.66).
▲
12
Strongly depleted amphibolites Depleted amphibolites
Middle crust 10
Enriched amphibolites
8
6
4
Crusta contaml trend ination
in the diagram of Figure 6 or in any other diagram using highly incompatible element-element or incompatible element ratios, thus confirming the OIB affinity. These two different trends are also seen through the variation in Nb/Y ratios relative to Zr/Y and Th/Yb versus Ta/Yb (Fig. 7A and B). Thus, the enriched amphibolites (Group C) plot within the mantle array similarly to OIB (see also Table 3), whereas the depleted amphibolites (Group B) show crustal contamination trends from an N-MORB–like depleted source owing to the relatively low Ta and Nb values of the crust. A discrimination factor based on the excess or deficiency in Nb has been used for identifying the mantle source (Fitton et al., 1997). This factor is expressed as:
100
Figure 5. Th vs. U diagram for the amphibolites studied, showing a linear trend for the enriched amphibolites. N-MORB, primitive mantle (PM), and OIB values after Sun and McDonough (1989).
Low-Ti tholeiite s from Bohemia
Figure 4. Trace element discrimination diagram after Meschede (1986). VAB—volcanic arc basalts; WPA—within-plate alkali basalts; WPT—within-plate tholeiites. Shaded area: data after Barbero and Villaseca (2000). Symbols as in Figure 5.
10
U (ppm)
Th/Ta
Zr/4
1
▲
2
OIB
ce Plume influen ) nd tre d (Icelan
N-MORB
0 0
200
400
600
Zr (ppm) Figure 6. Comparison of different trends in Central Iberian amphibolites using the Zr vs. Th/Ta diagram. Low-Ti Bohemian trend: thin arrow—crustal contamination trend: thick arrow—Iceland plume basalts: dashed arrow—data from Thompson et al. (1982, 1986) and Floyd et al. (2000). Middle crust: Rudnick and Taylor (1987), N-MORB and OIB values: Sun and McDonough (1989).
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López-Plaza et al.
A
B 10
Nb/Y
W PB
10
Sho OIB OIB
XMC X LC
IA UB
0.1
IA
Al k ns
Th/Yb
PM
Ca
1
EM
Thol
Th ol Tr a
X UC
1
0.1 N-MORB
DM
RB
LB
M
0 1
10
20
Ta/Yb
O
Zr/Y 0.01
0
0.1
1
10
Figure 7. (A) Nb/Y vs. Zr/Y diagram for the studied amphibolites. The enriched-amphibolite trend fits well with the Icelandic array, probably involving a primitive mantle component as the depleted end-member, whereas the depleted-amphibolite trend displays lower values of Nb/Y ratio, suggesting a crustal contamination (middle crust?). Lower crust (LC), middle crust (MC), and upper crust (UC) values after Rudnick and Fountain (1995). UBIA—upper bound of the Icelandic array; LBIA—lower bound of the Icelandic array (after Fitton et al., 1997). N-MORB, Primitive mantle (PM), and OIB values after Sun and McDonough (1989). (B) Th/Yb vs. Ta/Yb diagram after Pearce (1982). MORB and within-plate basalts (WPB) are subdivided into tholeiitic (Thol), transitional (Trans), and alkaline (Alk) varieties. Volcanic arc basalts are subdivided into tholeiitic (Thol), calc-alkaline (Ca), and shoshonitic (Sho) varieties. DM—depleted mantle; EM—enriched mantle. Symbols as in Figure 6.
Sample: Type: 87
86
Rb/ Sr 86 Sr/ Sr 87 86 ( Sr/ Sr)546 147 144 Sm/ Nd 143 144 Nd/ Nd 87
εNd 18 δ O‰ 546
TABLE 4. RADIOGENIC AND STABLE ISOTOPE DATA FOR SELECTED SAMPLES BR-15 DT-134 DT-135 DT-139B DT-142 DT-143 DT-136 DT-115 SDA DA DA DA EA EA EA EA 0.0611 0.707345 0.706869 0.2926 0.513292 +6.1 +4.0
0.0689 0.706582 0.706046 0.1927 0.512297 –6.4 n.d.
0.4823 0.709849 0.706095 0.1996 0.512793 +2.8 n.d.
0.2185 1.1705 0.9994 0.0798 0.711997 0.7161 0.710486 0.705547 0.710296 0.706989 0.702707 0.704926 0.1749 0.11 0.1727 0.136 0.512591 0.512661 0.512864 0.51266 +0.6 +6.5 +6.1 +4.7 +5.4 +8.9 +9.1 n.d.
CAS-5 EA
0.0459 0.1405 0.704448 0.70665 0.704091 0.705556 0.1543 0.1277 0.512713 0.512799 +4.4 +7.9 n.d. n.d.
Note: SDA—strongly depleted amphibolites; DA—depleted amphibolites; EA—enriched amphibolites.
The Nd and Sr isotope data (Table 4) also discriminate both types of amphibolites. The depleted amphibolites (Group B) show a wide Nd and Sr isotope range, including one negative εNd sample, in agreement with crustal contamination. The enriched amphibolites (Group C) display positive and high εNd values (εNd (546) = 7.9–4.4), suggesting a significant contribution from a depleted mantle component. A broad Sr radiogenic range (87Sr/86Sr (546) from 0.7027 to 0.7069) leads to a rough subhorizontal trend in the Sr-Nd isotopes (Fig. 8). Nevertheless, three samples plot within the mantle array and show a negative-slope trend. The rest of them display a certain departure above the mantle array trend and at the same time they have the highest values of Nd and Sr isotopes.
DISCUSSION Assimilation Fractional Crystallization (AFC) Processes for the Depleted Amphibolites The magnesium number appears to be a good indicator of magmatic processes owing to the good correlation with REE fractionation factors for Group B (Fig. 9A and B). One of the strongly depleted samples (Group A, BR-15) appears to be colinear with Group B samples, plotting at one end of the linear trend. This colinearity suggests that the strongly depleted sample might be a parental magma of the depleted amphibolites, involving a possible fractional crystallization process. Their relatively low
Contrasting mantle sources and processes involved in a peri-Gondwanan terrane 10
A
ntle ar r ay
6
4 Ma
8
307
3
(La/Sm)CN
εNd(546)
4 2 0 -2 -4
2
1
-6 -8 -10 0.701
(Sr/Sr)546 0.703
0.705
0.707
0.709
0 0.2
0.711
0.3
0.4
0.5
0.6
0.7
0.8
0.6
0.7
0.8
mgN
Figure 8. Sr vs. Nd isotope diagram for the amphibolites from the Central Iberian Zone. The reference age (546 Ma) was taken from Zeck et al. (2004), obtained for possible contemporary felsic gneisses. Symbols as in Figure 6.
4
B
(Gd/Yb)CN
3
Cr and Ni contents (Ni = 145 ppm; Cr = 439 ppm) tends to be a common feature of the continental sub-alkalic magmas (i.e., Wilson, 1989) and could be explained as being caused by the actual characteristics of the source, although late processes might have also disturbed their concentrations. To test the mineral association that has been presumably fractionated, a Pearce major element diagram (Pearce, 1968) was used (Fig. 10). On such a plot, a linear trend with a slope of 1.0 is seen for both the strongly depleted and depleted amphibolites, which suggests the plagioclase-clinopyroxene-olivine association as well as the lack of significant orthopyroxene and Fe-Ti oxide fractionation. Nevertheless, the broad ranges in Nd isotopes suggest that fractional crystallization is not the only process. Taking into account Sr and Nd isotopes along with trace element variation, two possibilities, which are not mutually exclusive, are suggested. On combining the variation in trace elements with respect to mgN and radiogenic Sr, it is possible to infer a probable wallrock alteration process, because the least-differentiated sample (higher mgN) shows the lowest εNd values and the highest crustal trace element affinity, such as a higher Zr content. Apart from the former possibility, an AFC process appears to be involved, because increasing incompatible element ratios versus any petrogenetic indicator of differentiation (i.e., Zr; Fig. 6) is a common geochemical feature of an AFC process. AFC modeling was approached using a metapelitic granulite from the lower continental crust in central Spain as contaminant (Table 5). Major element–based modeling was performed to
2
1
0 0.2
0.3
0.4
0.5
mgN Figure 9. Variation in (A) LREE and (B) HREE fractionation factors with respect to the magnesium number for the amphibolites studied. Arrows indicate differentiation trends for depleted amphibolites. Symbols as in Figure 6. Chondrite-normalizing values after Nakamura (1974).
establish the cumulate minerals involved, as well as the residual liquid, contaminant, and the assimilation-crystallization rate r (Table 5). Following this step, trace element AFC modeling was carried out using the equations of DePaolo (1981), and isotopic data were also included to check the validity of the AFC process. A very low assimilation-crystallization rate (r = 0.06) was
308
López-Plaza et al. 800
although Sr isotope–based modeling cannot be confirmed because of Sr mobility, in agreement with the scattered plots of immobile trace elements and mgN versus Sr isotope variation (not included).
Strongly depleted amphibolites
700
Depleted amphibolites 600
F′/K
500
Enriched Amphibolites and the Role Played by Mixing Processes
400 f(x) = 1.02x - 12.5 2 R = 0.99
300 200 100 0 0
200
400
600
800
Si/K Figure 10. Pearce element ratio diagram of F′/K vs. Si/K for depleted amphibolites and strongly depleted amphibolites. One of the latter (BR-15) was used as parental magma. F′ = 0.5(Fe + Mg) + 2Ca + 3Na. Symbols as in Figure 1.
obtained, and the recalculated percentages of cumulus minerals are plagioclase, 57.4; olivine, 20.5; and clinopyroxene, 22.1. The parental magma (the strongly depleted BR-15 sample) plots closest to the calculated solid residuum (Fig. 11A and B), which is a common feature for parental magmas of continental tholeiitic basalts that have undergone substantial low-pressure crystal fractionation (Cox, 1993; Best and Christiansen, 2001). Here each sample is treated as a magma pulse that was segregated separately from the primitive melts represented by BR-15. Concerning trace elements, in general AFC modeling shows a good fit between the measured (sample 139B) and calculated values for r = 0.06, according to estimates after major element modeling. Taking the latter into account, LREE and heavy REE (HREE) as well as some interelemental ratios, such as Zr/Nb and La/Yb (Fig. 12A–C), show an excellent fit. AFC modeling using Nd isotopic data confirms the previous results (Fig. 12D),
In general, the bivariate plots of two highly incompatible elements for the enriched amphibolites (Group C) show good correlations, as described above. In particular, the Th versus U diagram (Fig. 5) defines a straight line that intersects at the origin, indicating either similar bulk distribution coefficients between Th and U or a binary mixing process involving two end-members with similar Th/U ratios, as has been stated for alkaline volcanic rocks from southeast Spain (López-Ruiz and Cebriá, 1990). The lack of correlation between the incompatible elements or incompatible element ratios versus mgN (Fig. 9A and B) rules out any significant role of fractional crystallization and partial melting processes, favoring a mixing process in agreement with the conclusions inferred from Zr-Nb systematics (Kamber and Collerson, 2000). A linear correlation between Zr and Nb is not expected from variable degrees of melting, because the concentration of Nb is much more sensitive to the degree of melting than that of Zr, and the Zr-Nb systematics of OIB require binary mixing of two separate melts, as supported by regional studies (Weaver, 1991; Farley et al., 1992, among others). The depleted end-member, similar to MORB, appears to be common to all OIB, whereas in view of the highly variable Zr/Nb ratio the enriched end-member could reflect intramantle processes of fertilization. If these conclusions are applied, the Th-U correlation would imply a kind of mixing involving two end-members, both having similar and near-chondritic Th/U ratios (~3.6) (Table 3). The average values of Th/U for OIB (3.9) and the enriched amphibolites (3.3) are not consistent with a single N-MORB– like depleted-mantle component (~2.5) (Table 3). Therefore, a primitive mantle contribution having a near-chondritic Th/U ratio
TABLE 5. MAJOR ELEMENT–BASED LEAST-SQUARES MODELING FOR DEPLETED AMPHIBOLITES Co (BR15) SiO2 TiO2 Al2O3 FeO MnO MgO CaO K2O P2O5 X
49.25 0.68 18.17 8.59 0.16 10.02 13.05 0.05 0.03
Clinopyroxine 52.48 0.72 2.47 6.33 0.2 16.63 21.16 0.0 0.0 0.153
Olivine 38.89 0.0 0.0 20.95 0.0 39.79 0.37 0.0 0.0 0.142
Plagioclase
Contaminant (Metapelite U-96)
CL (DT 139B)
CL (calculated)
48.81 0.1 34.12 0.45 0.0 0.08 16.36 0.08 0.0 0.398
57.15 1.33 25.68 9.29 0.06 2.38 0.43 3.55 0.14 0.042
53.46 2.13 15.14 13.84 0.24 5.42 9.28 0.24 0.27 0.34
53.51 1.69 15.16 13.95 0.37 5.4 9.34 0.48 0.11
Note: Estimated assimilation-crystallization rate is r = 0.06; sum of the squared residuals is Σr = 0.31. C0— parental magma; CL—residual liquid; X—molar fraction of cumulated minerals and contaminant contribution. Sources: Mineral composition after López-Moro (2000); contaminant after Villaseca et al. (1999). 2
Plg 30 Contrasting mantle sources and processes involved in a peri-Gondwanan terrane
20
B
18
FRA
NAL
0
CRY
STA
LLIZ
ATIO
N
4
X
10
20
MgO
30
40
25
Al2O3
MELTING MANTLE PA RTIAL
Al2O3
6
Ol
Cpx 0
CTIO
14
8
10
Plg
16
10
20
30
BR15
12
Al2O3
A Low-pressure bulk mineral residuum
309
20 BR15
Ol
15
Lherzolite
Harzburgite
2
Cpx
TiO2
Dunite
10
0 0
0.5
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(~4) appears to be more plausible as a mixed depleted source. Similar conclusions can be drawn from other highly incompatible element ratios, such as Th/Ta, Th/Nb, and Th/La (Table 3). Thus, the ratios of both OIB and enriched amphibolites (Group C) show intermediate values between those for primitive mantle and depleted mantle. It is important to note that those ratios of primordial mantle and N-MORB with similar values, such as La/ Nb, result in a relative constancy for the corresponding ratios of the enriched amphibolites (Table 3). A key point in explaining the correlations among incompatible elements is provided by the Sr isotope variation. Sr radiogenic values correlate well with incompatible elements, resulting in linear trends that can be accounted for by mixing processes (Fig. 13). In contrast, few incompatible elements, such as Ti, correlate with Nd isotope values, indicating a decoupling between the Sr and Nd isotopic systems. The small variation in Nd isotope values must be explained in terms of source differences involving primitive mantle and NMORB–like components. The relatively narrow range of Nd isotope values (εNd(546), 6.08–4.4) for the enriched samples plotted within the mantle array (Fig. 8) indicates a stronger contribution of the depleted mantle component, in the 50–70% range, according to a simple mixing modeling. The broad range of Sr isotopes for samples departed from the mantle array can be explained as follows: 1. The correlations of the Sr isotope values versus some incompatible trace and some major elements (Fig. 13) indicate linear and hyperbolic trends with respect to the element and its reciprocal, supporting a mixing process. A correlation of Nd isotope values also would have been expected above solidus temperature, which is not the case. Thus, a fluid interaction appears to have been involved. 2. The good correlation of (87Sr/86Sr)546 values versus P (Fig. 13) is related to the presence of apatite. This mineral has been semiquantitatively explored by cathodoluminescence, affording a pale yellow luminescence as a consequence of high Mn and low REE contents (Wenzel and
Ramseyer, 1992). Taking into account the scarcity of Mn in apatite from mantle-derived magmas (KdapatMn = 0.13; Paster et al., 1974), a crustal contribution can be inferred from the yellow apatite crystals of enriched amphibolites (Group C). 3. High modal apatite contents should account for an increase in the Sm/Nd ratio and hence higher εNd values for the enriched amphibolites departed from the mantle array. The incorporation of Sm into the new apatite crystals might have occurred more readily than that of Nd, owing to its smaller ionic radius, resulting in an accelerated whole-rock 143Nd/144Nd evolution with time. Moreover, this apatite enrichment does not seem to have affected either Th/U ratios, in agreement with the apatite fractionation patterns determined for anatectic models (Bea et al., 1994), or other incompatible element ratios, such as Nb/Y, resulting in a scattered plot on ΔNb parameter versus (Sr/Sr)546 diagram (not included). 4. Some high delta oxygen values (δO18, +8.9 to +9.1) determined for enriched amphibolites (Group C) provide another indication of crustal contribution. In conclusion, ruling out a significant AFC process, opensystem behavior seems to have occurred during subsolidus evolution, resulting in a selective resetting for both the oxygen and Sr systems, whereas the Nd system reflects either a mixed source or an artifact, as may be deduced from the mantle-array samples or the departed ones, respectively. It is difficult to pinpoint the origin of such a crustal participation, but one possibility could be related to Variscan overprinting, through which a mobilization of phosphorous and other mobile elements could have occurred during anatectic events. Alkaline affinity is revealed by the mantle array samples having an OIB-like Nd isotopic signature as well as an incompatible-element enrichment that has scarcely been modified by any crustal contribution. Early Paleozoic alkaline magmas of an allochthonous unit have been reported in the northwest Iberia, emphasizing the role played by mantle fluids in an extremely thinned crust (Montero et al.,
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Contrasting mantle sources and processes involved in a peri-Gondwanan terrane 1998, in an allochthonous unit). In contrast, these sectors of the Central Iberian Zone, such as the area studied here, show abundant Cambro-Ordovician peraluminous felsic gneisses (i.e., the Ollo de Sapo gneiss; see Ortega and Gil-Ibarguchi, 1990), witnessing an old crustal anatexis. In turn these felsic gneisses underwent Variscan anatexis, resulting in voluminous S-type batholiths (Castro et al., 2000). Hence, the crustal contribution during subsolidus evolution of the mantle-derived magmas is consistent with all of the extensional anatectic events.
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The composition of the amphibolites seems to indicate the same magma types (tholeiitic and alkaline) and the same mantle-crust interplay throughout the whole autochthonous terrane of Central Iberian Zone. This probable homogeneity can be extended to other continental parts of the Gondwana margin, such as the central European Variscan Massifs (Floyd et al., 2000), suggesting a common evolution during the late NeoproterozoicCambro-Ordovician extension between Avalonia and its eastern continuation (Cadomia, Intra-Alpine terrane), in agreement with von Raumer et al. (2002).
CONCLUSIONS ACKNOWLEDGMENTS On the basis of trace element geochemistry, three types of amphibolites have been established, of which the enriched group (Group C) shows the highest εNd values. Strongly depleted and depleted amphibolites groups (Groups A and B) define a crustal contamination trend similar to the supposedly coeval mafic rocks of the Bohemian “low-Ti metatholeiites” and also similar to that of the Evora Massif in the Ossa-Morena Zone, whereas the enriched amphibolites display a “plume-influence” or OIBlike affinity trend similar to the early Paleozoic Bohemian alkali basalts. The excess or deficiency in Nb with respect to the Icelandic trend was used here for discriminating both crustal and OIB-like trends. A combination of Sr and Nd systematics and major element and incompatible element variations permit the identification of two major petrogenetic processes: 1. Magmatic differentiation processes for the depleted amphibolites can be inferred from a regular decrease in the magnesium number. The strongly depleted amphibolites could have been the parent magma, which underwent a low-pressure fractional crystallization process, removing plagioclase, clinopyroxene, and olivine. An AFC process and wall contamination are likely to have occurred, resulting in a broad range of Nd isotope values. 2. A mixing mantle array trend is reflected by some enriched amphibolites with high Nd and low Sr isotope signatures, which strongly suggests an important contribution of a depleted-mantle component. Other enriched amphibolites show high Sr and high O isotope signatures, indicating crustal contamination. In contrast, higher Nd isotope values for the latter amphibolites do not seem to reflect a crustal effect. This paradox can be accounted for by invoking open-system behavior during subsolidus evolution, resulting in an increase in phosphorous and hence abundant modal Mn-rich apatite crystals. Presumably, the presence of apatite could have brought about an increase in the bulk rock Sm/Nd ratio, giving rise to an abnormally high Nd isotope signature. The tholeiitic affinity for the depleted amphibolites as well as the alkaline affinity for some enriched amphibolites is consistent with an extensional within-plate setting. Moreover, an intracontinental back-arc setting related to the late Cadomian orogeny cannot be completely discounted for the depleted ones.
Financial support for this research was provided by the projects from the Junta de Castilla y León SA 53/97 and SA024/01, as well as from the Spanish Ministerio de Educación y Ciencia (CGL2004-06808-C04-04/BTE). The authors are grateful to Brendan Murphy and an anonymous referee, whose comments on the first draft of this article led to substantial improvements in the text. They are also indebted to I. Armenteros for making a cathodoluminescence microscope available and to N. Skinner, who reviewed the English text. REFERENCES CITED Anders, E., and Grevesse, N., 1989, Abundance of the elements: Meteoritic and solar: Geochimica et Cosmochimica Acta, v. 53, p. 197–214, doi: 10.1016/0016-7037(89)90286-X. Bandrés, A., Eguiluz, 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. Barbero, L., and Villaseca, C., 2000, Eclogite facies relics in metabasites from the Sierra de Guadarrama (Spanish Central system): P-T estimations and implications for the Hercynian evolution: Mineralogical Magazine, v. 64, p. 815–836, doi: 10.1180/002646100549814. Bea, F., Pereira, M.D., and Stroh, A., 1994, Mineral/leucosome trace-element partitioning in a peraluminous migmatite (a laser ablation-ICPMS study): Chemical Geology, v. 117, p. 291–312, doi: 10.1016/00092541(94)90133-3. Beetsma, J., 1995, The late Proterozoic/Paleozoic and Hercynian crustal evolution of the Iberian Massif, N Portugal, as traced by geochemistry and Sr-Nd-Pb systematics of pre-Hercynian terrigenous sediments and Hercynian granitoids [Ph.D. thesis]: Amsterdam, Vrije University, 223 p. Bellido, F., Capote, R., Casquet, C., Fúster, J.M., Navidad, M., Peinado, M., and Villaseca, C., 1981, Características generales del cinturón hercínico en el sector oriental del Sistema Central Español: Cuadernos de Geología Ibérica, v. 7, p. 15–51. Best, M.G., and Christiansen, E.H., 2001, Igneous petrology: New York, Blackwell Science, 458 p. Borthwick, J., and Harmon, R.S., 1982, A note regarding ClF3 as an alternative to BrF5 for oxygen isotope analysis: Geochimica et Cosmochimica Acta, v. 46, p. 1665–1668, doi: 10.1016/0016-7037(82)90321-0. Capote, R., Casquet, C., Fernández Casals, M.J., Moreno, F., Navidad, M., Peinado, M., and Vegas, R., 1977, The Precambrian in the central part of the Iberian Massif: Estudios Geológicos, v. 33, p. 343–355. Castro, A., Corretgé, L.G., El-Biad, M., El-Hmidi, H., Fernández, C., and Patiño Douce, A.E., 2000, Experimental constraints on Hercynian anatexis in the Iberian Massif, Spain: Journal of Petrology, v. 41, no. 10, p. 1471–1488. Clayton, R.N., and Mayeda, T.K., 1963, The use of bromine pentafluoride in the extraction of oxygen from oxides and silicates for isotopic analysis: Geochimica et Cosmochimica Acta, v. 27, p. 43–52, doi: 10.1016/00167037(63)90071-1.
312
López-Plaza et al.
Cox, K.G., 1993, Continental magmatic underplating: Philosophical Transactions of the Royal Society (Ser. A), v. 342, p. 155–166. Crowley, Q.G., Floyd, P.A., Winchester, J.A., Franke, W., and Holland, J.G., 2000, Early Paleozoic rift-related magmatism in Variscan Europe: Fragmentation of the Armorican terrane assemblage: Terra Nova, v. 12, p. 171–180, doi: 10.1046/j.1365-3121.2000.00290.x. DePaolo, D.J., 1981, Trace element and isotopic effects of combined wallrock assimilation and fractional crystallization: Earth and Planetary Science Letters, v. 53, p. 189–202, doi: 10.1016/0012-821X(81)90153-9. Escuder, J., and Navidad, M., 1999, Magmatismo prehercínico relacionado con extensión cortical en el Domo Gneísico del Tormes (Macizo Ibérico): evidencias geoquímicas en ortogneises y ortoanfibolitas, in XV Reunión Geología Oeste Peninsular (International Meeting on Cadomian Orogens), Annual meeting of the International Geological Correlation Program Project 376, Badajoz, Spain, October, p. 115–124. Farias, P., Gallastegui, G., González Lodeiro, F., Marquínez, J., Martín Parra, L.M., Martínez Catalán, J.R., De Pablo Macía, J.G., and Rodríguez Fernández, R., 1987, Aportaciones al conocimiento de la litoestratigrafía y estructura de la Galicia Central, in Comunicaçoes IX Reunión Geología Oeste Peninsular: Memorias 1: Porto, Portugal, Universidade de Oporto, p. 411–431. Farley, K.A., Natland, J.H., and Craig, H., 1992, Binary mixing of enriched and undegassed (primitive) mantle components (He, Sr, Nd, Pb) in Samoan lavas: Earth and Planetary Science Letters, v. 111, p. 183–199, doi: 10.1016/0012-821X(92)90178-X. Fernández-Suárez, J., Gutiérrez-Alonso, G., Jenner, G.A., and Jackson, S.E., 1998, Geochronology and geochemistry of the Pola de Allande granitoids (northern Spain): Their bearing on the Cadomian-Avalonian evolution of northwest Iberia: Canadian Journal of Earth Sciences, v. 35, p. 1439– 1453, doi: 10.1139/cjes-35-12-1439. Ferreira, N., Castro, P., Pereira, E., Dias, G., and Miranda, A., 2000, Syntectonic plutonism and Variscan anatexis of a Cadomian event, in Dias, G., Noronha, F., and Ferreira, N., eds., Variscan plutonism in the Central Iberian zone, northern Portugal: Eurogranites 2000, field meeting, guidebook: Portugal, p. 155–172. Fitton, J.G., Saunders, A.D., Norry, M.J., Hardarson, B.S., and Taylor, R.N., 1997, Thermal and chemical structure of the Iceland plume: Earth and Planetary Science Letters, v. 153, p. 197–208, doi: 10.1016/S0012821X(97)00170-2. Floyd, P.A., Winchester, J.A., Ciesielczuk, J., Lewandoska, A., Szczepanski, J., and Turniak, K., 1996, Geochemistry of early Palaezoic amphibolites from the Orlica-Snieznik Dome, Bohemian Massif: Petrogenetic and palaeotectonic aspects: Geologische Rundschau, v. 85, p. 225–238. Floyd, P.A., Winchester, J.A., Seston, R., Kryza, R., and Crowley, Q.G., 2000, Review of geochemical variation in lower Palaeozoic metabasites from the NE Bohemian Massif; intracratonic rifting and plume-ridge interaction, in Franke, W., Haak, V., Oncken, O., and Tanner, D., eds., Orogenic processes: Quantification and modelling in the Variscan belt: London, Geological Society of London Special Publication 179, p. 155–174. Friedman, I., and Gleason, J.D., 1973, Notes on the bromine pentafluoride technique of oxygen extraction: Journal of Research of the U.S. Geological Survey, v. 1, p. 679–680. Gómez-Pugnaire, M.T., Azor, A., Fernández-Soler, J.M., and López SánchezVizcaíno, V., 2003, The amphibolites from the Ossa-Morena/Central Iberian Variscan suture (southwestern Iberian Massif): Geochemistry and tectonic interpretation: Lithos, v. 68, p. 23–42, doi: 10.1016/S00244937(03)00018-5. Gonzalo, J.C., López-Plaza, M., Domínguez Vadillo, A., and López Moro, F.-J., 1994, Las anatexitas del Domo del Tormes y su significado litoestratigráfico (Prov. de Salamanca y Zamora): Boletín Geológico y Minero de España, v. 105, p. 397–411. Gutiérrez-Alonso, G., Fernández-Suárez, J., Jeffries, T.E., Jenner, G.A., Tubrett, M.N., Cox, R., and Jackson, S.E., 2003, Terrane accretion and dispersal in the northern Gondwana margin: An Early Palaeozoic analogue of a longlived active margin: Tectonophysics, v. 365, p. 221–232, doi: 10.1016/ S0040-1951(03)00023-4. Kamber, B.S., and Collerson, K.D., 2000, Zr/Nb systematics of ocean island basalts reassessed—The case for binary mixing: Journal of Petrology, v. 41, p. 1007–1021, doi: 10.1093/petrology/41.7.1007. Lefort, J.P., 1989, Basement correlation across the North Atlantic: Berlin, Springer-Verlag, 148 p. Leterrier, J., and Noronha, F., 1998, Evidências de um plutonismo calcoalcalino Cadomiano e de um magmatismo tipo MORB no Complexo Metamórfico
da Foz do Douro (Porto): Actas do V Congreso Nacional de Geologia Instituto Geológico e Mineiro de Portugal, v. 84, p. B-146–B-153. López-Moro, F.-J., 2000, Las Rocas Plutónicas Calcoalcalinas y Shoshoníticas del Domo Varisco del Tormes (Centro-Oeste Español) [Ph.D. thesis]: Salamanca, Spain, University of Salamanca, 441 p. López-Moro, F.-J., and López-Plaza, M., 1992, Geología del basamento en el antiforme de Miranda do Douro (sector oeste de Zamora): Studia Geologica Salmanticensia, v. 28, p. 103–141. López-Plaza, M., and López-Moro, F.-J., 1999, Procesos de fusión en relación con el plutonismo calcoalcalino y shoshonítico del Domo Varisco del Tormes (Zona Centro-Ibérica): XV Reunión Geología Oeste Peninsular (International Meeting on Cadomian Orogens), Annual Meeting of the International Geological Correlation Program Project 376, Badajoz, Spain, October, p. 175–183. López-Ruiz, J., and Cebriá, J.M., 1990, Geoquímica de los Procesos Magmáticos: Madrid, Rueda, 168 p. Martínez, F.J., 1974, Estudio del área metamórfica y granítica de los Arribes del Duero (Provincias de Salamanca y Zamora): Cuadernos Geología de Oviedo, v. 7, p. 3–141. Martínez Catalán, J.R., Arenas, R., Díaz García, F., and Abati, J., 1999, Allochthonous units in the Variscan belt of NW Iberia. Terranes and accretionary history, in Sinha, A.K., ed., Basement tectonics 13: Amsterdam, Kluwer Academic, p. 65–84. Meschede, M., 1986, A method of discriminating between different types of mid-ocean ridge basalts and continental tholeiites with the Nb-Zr-Y diagram: Chemical Geology, v. 56, p. 207–218, doi: 10.1016/00092541(86)90004-5. Montero, P., Floor, P., and Corretgé, L.G., 1998, The accumulation of rare-earth and high-field-strength elements in peralkaline granitic rocks: The Galiñeiro orthogneissic complex, northwestern Spain: Canadian Mineralogy, v. 36, p. 683–700. Murphy, J.B., and Nance, R.D., 1989, Model for the evolution of the Avalonian-Cadomian belt: Geology, v. 17, p. 735–738, doi: 10.1130/00917613(1989)017<0735:MFTEOT>2.3.CO;2. Murphy, J.B., Strachan, R.A., Nance, R.D., Parker, K.D., and Fowler, M.B., 2000, Proto-Avalonia: A 1.2–1.0 Ga tectonothermal event and constraints for the evolution of Rodinia: Geology, v. 28, p. 1071–1074, doi: 10.1130/0091-7613(2000)028<1071:PAAGTE>2.3.CO;2. Murphy, J.B., Eguiluz, L., and Zulauf, G., 2002, Cadomian orogens, peri-Gondwanan correlatives and Laurentia-Baltica connections: Tectonophysics, v. 352, p. 1–9, doi: 10.1016/S0040-1951(02)00186-5. Nakamura, N., 1974, Determination of REE, Ba, Mg, Na and K in carbonaceous and ordinary chondrites: Geochimica et Cosmochimica Acta, v. 38, p. 757–775, doi: 10.1016/0016-7037(74)90149-5. Nance, R.D., and Murphy, J.B., 1994, Contrasting basement isotopic signatures and the palinspastic restoration of peripheral orogens: Example from the Neoproterozoic Avalonian-Cadomian belt: Geology, v. 22, p. 617–620, doi: 10.1130/0091-7613(1994)022<0617:CBISAT>2.3.CO;2. Ordóñez-Casado, B., 1998, Geochronological studies of the pre-Mesozoic basement of the Iberian Massif; The Ossa Morena zone and the allochthonous complexes within the Central Iberian zone [Ph.D. thesis]: Zurich, University of Zurich, 235 p. Ortega, L.A., and Gil-Ibarguchi, J.I., 1990, The genesis of the late Hercynian granitoids from Galicia (NW Spain: inferences of REE studies): Journal of Geology, v. 98, p. 189–211. Paster, T.P., Schauwecker, D.S., and Haskin, L.A., 1974, The behaviour of some trace elements during solidification of the Skaergaard layered series: Geochimica et Cosmochimica Acta, v. 38, p. 1549–1577, doi: 10.1016/00167037(74)90174-4. Pearce, J.A., 1982, Trace element characteristics of lavas from destructive plate boundaries, in Thorpe, R.S., ed., Andesites: Chichester, Wiley, p. 525–548. Pearce, T.H., 1968, A contribution to the theory of variation diagrams: Contributions to Mineralogy and Petrology, v. 19, p. 142–157, doi: 10.1007/ BF00635485. Peinado, M., López-Plaza, M., Franco, P., Rodríguez-Alonso, M.D., Carnicero, A., and Gonzalo, J.C., 2002, Amphibolites from Central Iberian zone (Spain): Geochemical discrimination and processes involved: Programm und Kurzfassungen. Geo 2002, Würzburg, Germany, October: Planet Erde: Vergangenheit, Entwicklung, Zukunft, p. 262. Pereira, M.F., Chichorro, M., Santos, J.F., Moita, P., and Silva, J.B., 2004, Geochemistry of lower Paleozoic anorogenic basic rocks from the Evora Massif (western Ossa-Morena zone, Portugal): Geogaceta, v. 35, p. 87–90.
Contrasting mantle sources and processes involved in a peri-Gondwanan terrane Quesada, C., 1991, Geological constraints on the Paleozoic tectonic evolution of tectonostratigraphic terranes in the Iberian Massif: Tectonophysics, v. 185, p. 225–245, doi: 10.1016/0040-1951(91)90446-Y. Quesada, C., Bellido, F., Dallmeyer, R.D., Gil-Ibarguchi, J.I., Oliveira, J.T., Pérez Estaún, A., Ribeiro, A., Robardet, M., and Silva, J.B., 1991, Terranes within the Iberian Massif: Correlations with west African sequences, in Dallmeyer, R.D., and Lecorché, J.P., eds., The west African orogens and circum-Atlantic correlations: Berlin, Springer-Verlag, p. 267–294. Rodríguez-Alonso, M.D., and Alonso Gavilán, G., eds., 1995, NeoproterozoicLower Cambrian in the central-western part of the Iberian Peninsula: Post-Conference Field Guide, XIII Geological Meeting on the West of the Iberian Peninsula, 120 p. Rodríguez-Alonso, M.D., Díez Balda, M.A., Perejón, A., Pieren, A., Liñán, E., López Díaz, F., Moreno, F., Gámez Vintaned, J.A., González Lodeiro, F., Martínez Poyatos, D., and Vegas, R., 2004a, La secuencia litoestratigráfica del Neoproterozoico-Cámbrico Inferior, in Vera, J.A., ed., Geología de España: Madrid, Sociedad Geológica de España–Instituto Geológico y Minero de España, p. 78–81. Rodríguez-Alonso, M.D., Peinado, M., López-Plaza, M., Franco, M.P., Carnicero, A., and Gonzalo, J.C., 2004b, Neoproterozoic-Cambrian synsedimentary magmatism in the Central Iberian zone (Spain): Geology, petrology and geodynamic significance: International Journal of Earth Sciences, v. 93, p. 897–920, doi: 10.1007/s00531-004-0425-4. Rudnick, R.L., and Fountain, D.M., 1995, Nature and composition of the continental crust: A lower crustal perspective: Reviews of Geophysics, v. 33, p. 267–310, doi: 10.1029/95RG01302. Rudnick, R.L., and Taylor, S.R., 1987, The composition and petrogenesis of the lower crust; A xenolith study: Journal of Geophysical Research, v. 92, p. 13,981–14,005. Sánchez-García, T., Bellido, F., and Quesada, C., 2003, Geodynamic setting and geochemical signatures of Cambrian-Ordovician rift-related igneous rocks (Ossa-Morena zone, SW Iberia): Tectonophysics, v. 365, p. 233– 255, doi: 10.1016/S0040-1951(03)00024-6. Skehan, J.V., and Rast, N., 1995, Late Proterozoic to Cambrian evolution of the Boston Avalon Terrane, in Hibbard, J.P., Van Staal, C.R., and Cawood, P.A. (eds), Current perspectives in the Appalachian-Caledonian Orogen: Geol. Assoc. Canada Spec. Paper, v. 41, p. 207–225. Späth, A., Le Roex, A., and Opiyo-Akech, N., 2001, Plume-lithosphere interaction and the origin of continental rift-related alkaline volcanism—The Chyulu Hills volcanic province, southern Kenya: Journal of Petrology, v. 42, p. 765–787, doi: 10.1093/petrology/42.4.765. Sun, S.S., and McDonough, W.F., 1989, Chemical and isotopic systematics of oceanic basalts: Implications for mantle composition and processes, in Saunders, A.D., and Norry, M.J., eds., Magmatism in the ocean basins: London, Geochemical Society of London Special Publication 42, p. 313–345.
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Taylor, S.R., and McLennan, S.M., 1985, The continental crust: Its composition and evolution: New York, Blackwell Science, 312 p. Thompson, R.C., Dickin, A.P., Gibson, I.L., and Morrison, M.A., 1982, Elemental fingerprints of isotopic contamination of Hebridean Palaeocene mantle-derived magmas by Archean sial: Contributions to Mineralogy and Petrology, v. 79, p. 159–168, doi: 10.1007/BF01132885. Thompson, R.N., Morrison, M.A., Hendry, G.L., and Parry, S.J., 1984, An assessment of the relative roles of crust and mantle in magma genesis: An elemental approach: Philosophical Transactions of the Royal Society of London (Ser. A), v. 310, p. 549–590. Thompson, R.N., Morrison, M.A., Dickin, A.P., Gibson, I.L., and Harmon, R.S., 1986, Two contrasting styles of interaction between basic magmas and continental crust in the British Tertiary volcanic province: Journal of Geophysical Research, v. 91, p. 5985–5997. Valverde Vaquero, P., and Dunning, G.R., 2000, New U-Pb ages for Early Ordovician magmatism in central Spain: Journal of the Geological Society of London, v. 157, p. 15–26. Villaseca, C., Downes, H., Pin, C., and Barbero, L., 1999, Nature and composition of the lower continental crust in central Spain and the granulite-granite linkage: Inferences from granulitic xenoliths: Journal of Petrology, v. 40, p. 1465–1496. von Raumer, J.F., Stampfli, G.M., Borel, G., and Bussy, F., 2002, Organization of pre-Variscan basement areas at the north-Gondwanan margin: International Journal of Earth Sciences, v. 91, p. 35–52, doi: 10.1007/ s005310100200. Weaver, B.L., 1991, Trace-element evidence for the origin of ocean-island basalts: Geology, v. 19, p. 123–126, doi: 10.1130/0091-7613(1991)019<0123: TEEFTO>2.3.CO;2. Wenzel, T., and Ramseyer, K., 1992, Mineralogical and mineral-chemical changes in a fractionation-dominated diorite-monzodiorite-monzonite sequence: Evidence from cathodoluminescence: European Journal of Mineralogy, v. 4, p. 1391–1399. Wilson, M., 1989, Igneous petrogenesis. A global tectonic approach: London, Unwin Hyman, 466 p. Winchester, J.A., and Floyd, P.A., 1977, Geochemical discrimination of different magma series and their differentiation products using immobile elements: Chemical Geology, v. 20, p. 325–343, doi: 10.1016/00092541(77)90057-2. Zeck, H.P., Wingate, M.T.D., Pooley, G.D., and Ugidos, J.M., 2004, A sequence of pan-African and Hercynian events recorded in zircons from an orthogneiss from the Hercynian belt of western central Iberia—An ion microprobe U-Pb study: Journal of Petrology, v. 45, no. 8, p. 1613–1629, doi: 10.1093/petrology/egh026. MANUSCRIPT ACCEPTED BY THE SOCIETY 3 OCTOBER 2006
Printed in the USA
Geological Society of America Special Paper 423 2007
Tectonic evolution of the upper allochthon of the Órdenes complex (northwestern Iberian Massif): Structural constraints to a polyorogenic peri-Gondwanan terrane Juan Gómez Barreiro* Department of Earth and Planetary Sciences, University of California, 305 McCone Hall, Berkeley, California 94720, USA José R. Martínez Catalán Departamento de Geología, Universidad de Salamanca, Plaza de los Caidos s/n, 37008 Salamanca, Spain Ricardo Arenas Pedro Castiñeiras Jacobo Abati Departamento de Petrología y Geoquímica, Universidad Complutense, José Antonio Novais 2, 28040 Madrid, Spain Florentino Díaz García Departamento de Geología, Universidad de Oviedo, Jesús Arias de Velasco s/n, 33005 Oviedo, Spain Jan R. Wijbrans Department of Isotope Geochemistry, Vrije Universiteit, De Boelelaan 1085, Amsterdam 1081 HV, The Netherlands
ABSTRACT The upper allochthon of northwest Iberia represents the most exotic terrane of this part of the European Variscan belt. Recent advances in the metamorphic petrology, structural geology, and geochronology of the upper allochthon in the Órdenes complex are integrated into a synthesis of its tectonic evolution, constraining the main tectonothermal events. Important aspects of this synthesis are (1) the interpretation of Cambro-Ordovician magmatism and earliest metamorphic event, as the result of drifting of a peri-Gondwanan terrane; (2) the subsequent shortening and crustal thickening of the terrane related to its subduction and accretion to Laurussia; (3) a younger cycle of shortening and extension resulting from convergence between Laurussia and Gondwana; and (4) the emplacement of this exotic terrane as the upper allochthon, together with underlying ophiolitic and basal allochthons, during the Laurussia-Gondwana collision. Implications derived from the well-established tectonothermal sequence are discussed in the context of Paleozoic paleogeography and geodynamics. The evolution *E-mail:
[email protected]. Gómez Barreiro, J., Martínez Catalán, J.R., Arenas, R., Castiñeiras, P., Abati, J., Díaz García, F., and Wijbrans, J.R., 2007, Tectonic evolution of the upper allochthon of the Órdenes complex (northwestern Iberian Massif): Structural constraints to a polyorogenic peri-Gondwanan terrane, 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. 315–332, doi: 10.1130/2007.2423(15). For permission to copy, contact
[email protected]. ©2007 Geological Society of America. All rights reserved.
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Gómez Barreiro et al. of this part of the belt is related first to the closure of the Tornquist Ocean, and later to that of the eastern branch of the Rheic Ocean. Furthermore, the relative paleopositions of the upper allochthon and the Iberian autochthon in northern Gondwana are discussed. Keywords: tectonics, polyorogenic terrane, Variscan orogeny, Gondwana
INTRODUCTION The reconstruction of Paleozoic continents, intervening oceans and peri-cratonic terranes relies to a great extent on detailed structural and petrologic analyses of exotic terranes occurring around suture zones, as well as on precise age determinations. The Variscan belt of central and western Europe offers the opportunity to study such types of terranes in a strip running from northwest Iberia to the Bohemian Massif through the French Armorican and Central massifs (Fig. 1). Ophiolitic units mark the closure of a Paleozoic ocean in several of these massifs. In northwest Iberia, allochthonous ophiolitic units highlight a rootless suture above which an exotic terrane, including high-P/high-T eclogites and granulites, occurs. This upper allochthon has been widely studied. It consists of metasediments and igneous felsic, mafic, and ultramafic rocks, with gabbroic and granitic protoliths dated at ca. 500 Ma (Peucat et al., 1990; Dallmeyer and Tucker, 1993; Schäfer et al., 1993; Abati et al., 1999) and also characterized by early Variscan (425– 390 Ma) metamorphism (Schäfer et al., 1993; Santos Zalduegui et al., 1996; Dallmeyer et al., 1997; Ordóñez Casado et al., 2001; Roger and Matte, 2005; Gómez Barreiro et al., 2006; FernándezSuárez et al., 2007). Units with similar lithologies and ages that have experienced high-P early Variscan metamorphism also exist in the French Central Massif (Ledru et al., 1994; Ploquin and Santallier, 1994; Santallier et al., 1994), the Vosges and the Black Forest (Franke, 2000), the Bohemian Massif (Franke, 2000; Franke and Zelazniewicz, 2000; Crowley et al., 2002), and also in the Alpine basement (Neubauer et al., 1999). Many of these units can be related to the upper Iberian allochthon, and all together might have formed a coherent terrane whose history we are trying to unravel. In this contribution, detailed tectonometamorphic studies in three units of the upper allochthon of northwest Spain (Abati, 2002; Castiñeiras, 2003; Gómez Barreiro, 2004), combined with U-Pb age determinations (Abati et al., 1999), 40Ar/39Ar dating of regional fabrics (Gómez Barreiro et al., 2006), and structural, thermobarometric and age data from the rest of the Spanish upper allochthon, are integrated in an evolutionary model for this exotic terrane widely represented in the European Variscides. GEOLOGICAL SETTING The allochthonous complexes of northwest Iberia (Fig. 2) form a stack of exotic terranes preserved as megaklippen in structural basins interspersed among migmatite and granite domes
that developed during gravitational collapse and extension in advanced stages of the Variscan orogeny (Martínez Catalán et al., 2007). In a way similar to what has been described for the Caledonian belt of Scandinavia (Gee and Sturt, 1985; Roberts and Gee, 1985; Stephens and Gee, 1989; Rey et al., 1997), upper, middle, and basal allochthons were thrust onto the Iberian autochthon (the Central Iberian zone) with an intervening parautochthonous thrust sheet between them (Farias et al., 1987; Ribeiro et al., 1990). Other exotic units of similar meaning have been described in the Ossa-Morena zone, in southwest Iberia (Fig. 1), by Leal et al. (1996), Araújo and Ribeiro (1997), Fonseca (1997), and Fonseca et al. (1999). There has been some debate concerning the permanence of the autochthon in Gondwana or its individualization as part of the Armorica plate. The Armorica hypothesis (Van der Voo, 1982, 1988) is based on paleomagnetism and has been questioned on the same grounds by Kent et al. (1984), Scotese (1984), and Hargraves et al. (1987). Paris and Robardet (1990) and Robardet et al. (1990) argued against a significant separation of Armorica from Gondwana, based on sedimentary and faunal evidence. Robardet (2002, 2003) concluded that only models involving a single Rheic Ocean between Gondwana and Laurentia are realistic on the basis of paleogeographic analysis. Moreover, U-Pb dating of detrital zircons in low-grade metasedimentary rocks of the Iberian autochthon ranging in age from the Neoproterozoic to the Early Devonian have yielded the same age populations, suggesting a common source area during the whole Paleozoic and militating against microplate separation (Fernández-Suárez et al., 2000, 2002a,b; Martínez Catalán et al., 2004). Turning back to the exotic terranes of northwest Iberia, the middle allochthon consists entirely of ophiolites. Two different origins and ages have been proven for the several existing ophiolitic units (Díaz García et al., 1999a; Pin et al., 2002; Arenas et al., 2004, 2007; Sánchez Martínez et al., 2007), representing different oceanic lithospheres formed at different stages inside the Rheic oceanic realm. Both the parautochthon and the basal allochthon are considered Gondwanan pieces that never abandoned the northern margin of the continent. Conversely, the upper allochthon clearly represents an exotic terrane, as demonstrated by the structurally underlying ophiolites. However, it is considered to have a Gondwanan origin because the oldest ages obtained from upper intercepts and inherited zircons from orthogneisses, paragneisses, and metabasites (between 2.7 and 1.8 Ga; Kuijper, 1980; Peucat et al., 1990; Dallmeyer and Tucker, 1993; Schäfer et al., 1993) are similar to those found in the orthogneisses of the basal allochthon
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Figure 1. Sketch showing the position of Iberia in relation with the Appalachian, Caledonian, and Variscan belts at the end of the Paleozoic convergence. The continents involved in the successive collisions and the main massifs of the European Variscides are also indicated. LBM— London-Brabant Massif; STA—Silesian terrane assemblage; stippled area—Avalonia. Modified from Martínez Catalán et al. (2002) and based on Neuman and Max (1989).
(1.8 Ga; Santos Zalduegui et al., 1995) and of the autochthon (Lancelot et al., 1985; Gebauer, 1993). And they are also similar to those of northern Africa, including the west African craton and the Saharan basement remobilized during the pan-African orogeny (Caby, 1989; Rocci et al., 1991; Avigad et al., 2003). These similarities point to a common Gondwanan basement for the upper and basal allochthons and for the Iberian autochthon. Moreover, graywackes from low-grade metasediments of the uppermost unit in the Órdenes complex have been investigated for detrital zircon ages, yielding three age populations of
2.5–2.4 Ga, 2.1–1.9 Ga, and 610–480 Ma (Fernández-Suárez et al., 2003), which record the major events in the west African craton of northern Gondwana and the surrounding pan-African belts (Caby, 1989; Rocci et al., 1991). Notably, the Neoproterozoic to Early Devonian sediments of the autochthon underlying the complexes have yielded detrital zircons with similar age populations and, in addition, a Mesoproterozoic (1.2–0.9 Ga) age cluster (Fernández-Suárez et al., 1999, 2000, 2002a,b; Martínez Catalán et al., 2004). These age data suggest that even if the upper allochthon is a terrane drifted away
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Figure 2. The Iberian Massif showing the subdivision in zones, based on Julivert et al. (1972) and Farias et al. (1987). Allochthonous complexes of northwest Iberia: B—Bragança; CO—Cabo Ortegal; M—Morais; MT—Malpica-Tui; O—Órdenes. Transcurrent shear zones: BCSZ—Badajoz-Córdoba shear zone; JPSZ—Juzbado-Penalva shear zone; MVSZ—Malpica-Vigo shear zone; PTSZ—Porto-Tomar shear zone.
from Gondwana and later emplaced on it, the site of derivation differs from that of emplacement. Late Cambrian to Early Ordovician magmatism is widespread in the upper and basal allochthons and also in the autochthon, being a little older in the upper allochthon (ca. 500 Ma; Dallmeyer and Tucker, 1993; Abati et al., 1999) than in the basal allochthon and the autochthon (490–470 Ma; Van Calsteren et al., 1979; García Garzón et al., 1981; Vialette et al., 1987; Gebauer, 1993; Santos Zalduegui et al., 1995; Valverde Vaquero and Dunning, 2000). That magmatism has calcalkaline and arc affinities in the autochthon (Ortega et al., 1996), in some upper allochthonous units and in many granitoids of the basal allochthon, but some granites and mafic rocks of the latter are alkaline to per-alkaline (Floor, 1966; Pin et al., 1992). The problem is to reconcile the tectonic stability registered by the Early Ordovician sediments of the autochthon, typical of a passive margin, the alkaline, rift-related magmatism of the basal allochthon, the widely accepted Early Paleozoic terrane dispersion in the peri-Gondwanan realm, and the calcalkaline and arc
affinities in the upper allochthon and the autochthon. Valverde Vaquero and Dunning (2000) have suggested that rifting was located in a back-arc setting behind a subduction zone. This hypothesis is supported by Stampfli and Borel (2002), Stampfli et al. (2002), Winchester et al. (2002), and von Raumer et al. (2003) in their reconstructions of the peri-Gondwanan terranes and implies that rolling back of the subducting slab of a wide ocean may have pulled the future upper allochthonous terrane apart from Gondwana, as the subduction of the southern Iapetus margin did with Avalonia. THE UPPER ALLOCHTHON IN THE ÓRDENES COMPLEX Different tectonic units can be distinguished in the upper allochthon, being separated from each other by faults, often extensional detachments (Martínez Catalán et al., 2002). These units can be grouped, according to their metamorphic imprint, into two different sets: (1) high-P/high-T units, and (2) intermediate-P
Tectonic evolution of the upper allochthon of the Órdenes complex units. Both sets exist in the Spanish allochthonous complexes of Cabo Ortegal and Órdenes (Fig. 2). The first set is better exposed in Cabo Ortegal, where very good samples of high-P/high-T fabrics have been preserved (Abalos et al., 1996; Abalos, 1997; Abalos and Aranguren, 1998). The second set shows its best outcrops in Órdenes (Martínez Catalán et al., 2002). The high-P/high-T units occupy a lower structural position below the intermediate-P units (Fig. 3), the contact being systematically of extensional character (Díaz García et al., 1999b; Castiñeiras, 2003; Gómez Barreiro, 2004; González Cuadra, 2005). Any indication of a possible suture between both sets is lacking, so that they are considered as parts of the same tectonostratigraphic terrane in spite of their different metamorphic evolutions. Actually, their metamorphic evolutions show a complementary history for both sets, thermobarically converging when approaching the transition zone, that is, the detachment separating them. These zones of contact may show a subsequent record of overprinting events and reworking (Gómez Barreiro, 2004). Our study focuses on the Órdenes complex, the largest (135 by 75 km) of the klippen containing exotic terranes in northwest Iberia. Detailed descriptions and interpretations of many of its units corresponding to the upper allochthon have been published elsewhere (Van Zuuren, 1969; Hubregtse, 1973; Díaz García, 1990; Abati et al., 1999, 2003; Díaz García et al., 1999b; Andonaegui et al., 2002; Arenas and Martínez Catalán, 2002). Figure 3 shows a simplified geological map of the complex and a schematic cross-section. For a more detailed map and structural description, the reader is referred to Martínez Catalán et al. (2002). Figure 4 depicts a cross-section representing lithologies of the upper allochthonous units together with protolith and metamorphic ages, whereas Figure 5 shows the same section with the maximum metamorphic grade reached by the different units and the P-T paths established for several of them. As can be seen in Figure 5, metamorphism increases from top to bottom, from greenschist to high-T eclogite facies, although the transition, far from being gradual, is accomplished via discrete extensional detachments. Furthermore, because of the lenticular rather than tabular character of many units, rocks with a great difference in their peak P-T conditions occur adjacent to one another. The fact that units with differences of more than 1.2 GPa in their pressure peak occur presently in a sheet ~10 km thick indicates that the original pile has been largely attenuated. The units of Betanzos, O Pino, Corredoiras, and Monte Castelo are of the intermediate-P type, whereas those of Fornás, Arinteiro, Belmil, Melide, and Sobrado correspond to the highP/high-T type. A brief description of each unit follows, given in descending order of what supposedly was their original structural position within the orogenic wedge. Protolith ages and depositional age constraints are given, when available, whereas metamorphic ages are presented and discussed later, together with the tectonometamorphic evolution. The Betanzos unit consists of a monotonous succession of slates and metagraywackes with turbiditic characteristics and also contains minor quartzites, conglomerates, and diabase dikes.
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Graywackes are feldspathic with a framework of quartz and fresh plagioclase. Rock fragments with vitric and microgranular textures are common in the polymictic conglomerates and coarsegrained graywackes, together with slates, cherts, and clasts of bipyramidal quartz. The abundance of volcanic components suggests a volcanic environment. The metagraywackes have been investigated for detrital zircon ages, yielding three age populations of 2.5–2.4 Ga, 2.1–1.9 Ga, and 610–480 Ma (FernándezSuárez et al., 2003), which establishes a maximum depositional age at the Early–Middle Ordovician transition. The sediments were affected by greenschist-facies metamorphism and by early recumbent and late upright folds (Matte and Capdevila, 1978). O Pino unit is a thick pile of mesozonal monotonous schists and paragneisses intruded by relatively small bodies of gabbro and granitoids. It overlies the huge meta-igneous massifs that form the two underlying units. This is a Barrovian-like pile, with metamorphic zones ranging from almandine to sillimanite (Abati, 2002; Castiñeiras, 2003). Monazites from semipelitic paragneisses of the sillimanite zone have yielded ages of 493 ± 1.3 and 496 ± 3 Ma (Abati et al., 1999). The Corredoiras unit is a coarse-grained, massive metagranodiorite and subordinate tonalite, whose crystallization is dated at 500 ± 2 Ma (U-Pb dating on zircons; Abati et al., 1999). The massif is variably gneissified and kilometer-scale xenoliths of metasediments occur inside the orthogneisses. The xenoliths appear frequently migmatized, but also locally granulitized, with a paragenesis indicative of the intermediate-P granulite facies (González Cuadra, 2005). This high-grade, intermediate-P metamorphism affected the whole unit. The Monte Castelo unit is a two-pyroxene gabbro of tholeiitic character, similar in composition to modern island arc basalts (Andonaegui et al., 2002). The presence of olivine and the common ophitic textures point to a relatively shallow emplacement. U-Pb analyses on zircons yielded an age of 499 ± 2 Ma for crystallization of the protolith (Abati et al., 1999). Scarce metapelitic intercalations and a kilometer-scale shear zone inside the mostly undeformed gabbro developed intermediate-P granulite-facies parageneses. This unit is considered the lowermost intermediate-P unit and, although it reached granulite facies, its metamorphic gradient was never of the high-P type (Abati, 2002; Abati et al., 2003). The Fornás and Arinteiro units are similar, the latter being a horse of the former sandwiched between thrust faults. They consist of heterogeneous and well-foliated amphibolites. Van Zuuren (1969) described preserved igneous textures, indicating a gabbroic origin, and also granulite-facies relics, which have been confirmed by Gómez Barreiro (2004). The dominant amphibolite-facies foliation developed subsequently, during the exhumation, in an extensional shear zone that drove the unit into contact with the overlying mesozonal schists and paragneisses of the O Pino unit. This shear zone, known as the Fornás detachment, was subsequently affected by recumbent folds, repeated by thrusts, and back-folded during the motion of the younger Bembibre–Pico Sacro extensional detachment (Figs. 3 and 4). Fornás
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Iberian M
VARISCAN GRANITOIDS Migmatites Two-mica aluminous granites
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ÓRDENES COMPLEX UPPER ALLOCHTHON Intermediate-P units BETANZOS UNIT Metagraywackes, slates, phyllites, conglomerates, and quartzites
Betanzos
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Carballo PCD
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Monte Castelo Bazar
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High-P / high-T units FORNÁS, SOBRADO, BELMIL AND MELIDE UNITS Paragneisses, amphibolites, basic granulites, eclogites and ultramafic rocks OPHIOLITIC ALLOCHTHON CAREÓN AND BAZAR UNITS Flasergabbros, amphibolites and ultramafic rocks VILA DE CRUCES UNIT Greenschist-facies metabasites, phyllites and schists
Santa Comba 43°
Ordes (Órdenes) PCD
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Ponte Carreira
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BASAL ALLOCHTHON AGUALADA, SANTIAGO, FORCAREI AND LALÍN UNITS Schists, albitic schists, paragneisses, amphibolites and felsic orthogneisses
LFT
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PARAUTOCHTHON SCHISTOSE DOMAIN Schists, carbonaceous schists, and metaquartzites
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Belmil unit Vila de Cruces unit PCD
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Careón unit Melide unit
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Negreira granodiorite Santiago unit
SE Corredoiras unit
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Figure 3. Geological map and cross-section (A-A′) of the Órdenes complex (northwest Spain) showing the allochthonous units and the main faults. BPSD—Bembibre–Pico Sacro detachment; CD—Corredoiras detachment; CMD—Campo Marzo detachment; FD—Fornás detachment; PCD—Ponte Carreira detachment; LFT—Lalín-Forcarei thrust.
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CD
Figure 4. Distribution of the main lithologies of the upper allochthon in a cross-section of the Órdenes complex and published isotopic ages of igneous protoliths (black box) and regional fabrics (white box), with indications of the corresponding metamorphic stage. Numbers in parentheses refer to references: 1—Abati et al. (1999); 2—Dallmeyer et al. (1997); 3—Fernández-Suárez et al. (2007); 4—Gómez Barreiro (2004).
Figure 5. Distribution of the maximum metamorphic grade reached by the units of the upper allochthon in a cross-section of the Órdenes complex and published pressure-temperature paths for several units. Dashed lines in P-T diagrams represent aluminum silicate phase diagram after Holdaway (1971). Numbers in parentheses refer to references: 1—Castiñeiras (2003); 2—Gómez Barreiro (2004); 3—Abati et al. (2003); 4—Arenas and Martínez Catalán (2002); 5—Díaz García et al. (1999b).
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is the uppermost high-P/high-T unit and the one in which the high-pressure parageneses (Van Zuuren, 1969; Gómez Barreiro, 2004) have been less preserved by the subsequent amphibolitefacies evolution. The Belmil, Melide, and Sobrado units are very similar, and also equivalent to high-P/high-T units in the Cabo Ortegal complex, where they have been widely studied. They consist of paragneisses and mafic and ultramafic meta-igneous rocks. Mafic rocks derive from gabbros of tholeiitic composition whose geochemical signature has been compared to mid-ocean ridge basalt (MORB; Gil-Ibarguchi et al., 1990) and related with continental rifting (Galán and Marcos, 1997). Conversely, geochemical studies in the ultramafic rocks of Cabo Ortegal point to a suprasubduction origin for these pyroxenite-rich peridotites at ca. 500 Ma (Santos et al., 2002). The mafic igneous protoliths have been dated at 520–480 Ma (Peucat et al., 1990; Ordóñez Casado et al., 2001; Fernández-Suárez et al., 2007). Field and map relationships show that they intruded the sedimentary precursor of the paragneisses. These have yielded U-Pb sensitive high-resolution ion microprobe (SHRIMP) core ages clustering around 2.0–2.2 Ga and 540 Ma, with the youngest detrital zircon dated at 507 Ma (Schäfer et al., 1993), thus establishing a Late Cambrian maximum depositional age. The metabasites include garnet-clinopyroxene granulites and eclogites, retrograded to the amphibolite facies (Vogel, 1967; Hubregtse, 1973). In Sobrado, gabbros occur in several stages of transformation, from practically undeformed and scarcely affected by the metamorphism to coronitic metagabbros and high-P granulites. In the less-deformed gabbros, subophitic and diabase textures have been preserved, indicating an emplacement at relatively shallow crustal levels (Arenas and Martínez Catalán, 2002). The tectonothermal evolution includes an old high-P granulite to eclogite facies metamorphism, followed by decompression and partial melting and then, successively, by a penetrative mylonitization in the amphibolite facies, recumbent folding, and thrusting in the greenschist facies (Vogel, 1967; Gil-Ibarguchi et al., 1990; Arenas, 1991; Girardeau and Gil-Ibarguchi, 1991; Mendia Aranguren, 2000; Marcos et al., 2002). TECTONOMETAMORPHIC EVOLUTION The following evolution has been deduced from detailed studies of the O Pino, Monte Castelo, and Fornás units, and this description focuses on the extensional detachments separating them. The combined investigation of large structures and thin sections with thermobarometric and age determinations have revealed complex P-T paths characterized by two loops, the first of which is counterclockwise for the intermediate-P units (Fig. 5). This fact provides the guideline for the following description, which splits the metamorphic and deformational history into six correlative tectonothermal stages contemporaneous with plate convergence, and a previous one related to terrane individualization.
D0-M0: A Thermal Event Related to Magmatism The oldest structural and metamorphic relicts in the O Pino unit are mesoscopic veins with andalusite and quartz that show prekinematic relationships with the first identified tectonic foliation and are preserved in some pelitic gneisses. These veins have been related to low-P heating and linked to the intrusion of huge plutonic massifs, such as those partially preserved in the Monte Castelo and Corredoiras units. Consequently, the oldest loop of the metamorphic P-T path is drawn starting with isobaric heating related to intrusion of massive gabbro and granodiorite bodies (Abati et al., 2003; Castiñeiras, 2003). Pressure and maximum temperature reached in this stage in the Monte Castelo unit have been estimated by Abati et al. (2003) to be 4–6 kbar and 800 °C. Monazite included in biotite in large enclaves of pelitic composition yielded ages of 498 ± 2 Ma in the Monte Castelo gabbro and 493 ± 2 and 484 ± 2 Ma in the Corredoiras orthogneiss (Abati et al., 1999), close to the age of the host plutonic rocks (499 ± 2 and 500 ± 2 Ma, respectively). The arc affinities of the Monte Castelo gabbro (Andonaegui et al., 2002) together with the immature character of the graywackes of the Betanzos unit and their high content in volcanogenic fragments suggest a relationship of isobaric heating with magmatic underplating and massive gabbro and granodiorite-tonalite intrusion in relatively shallow crustal levels of a magmatic arc. Contemporaneous partial melting of metasediments is documented by a granitoid rich in pelitic enclaves occurring at the roof of the Monte Castelo gabbro and dated at 500 ± 2 Ma (Abati et al., 1999). The previous data pertain to the intermediate-P units. Although in the high-P/high-T units the mafic protoliths have the same Cambro-Ordovician ages, their geochemistry points rather to continental rifting (Galán and Marcos, 1997). This rifting is not in conflict with the arc affinities of the overlying units, as both can be integrated in the picture of a terrane drifting away from Gondwana and pulled by slab roll-back of a subduction zone that provides the arc-related magmatism (Fig. 6). D1-M1: First Burial by Crustal Thickening and Subduction The first thermal event D0-M0 was followed by pressurization related to the accretion of the upper allochthon to an active orogenic wedge (Martínez Catalán et al., 1997, 1999). Both sets of the upper allochthon reached metamorphic peak conditions during this event. In the intermediate-P units, the regional foliation (S1) developed, varying from slaty cleavage at the uppermost levels to gneissic layering at depth as metamorphic grade increased (Díaz García, 1990; Abati, 2002; Castiñeiras, 2003; Gómez Barreiro, 2004). Petrographic evidence suggests that rocks were pressurized following an isothermal path according to Abati et al. (2003) and Castiñeiras (2003), reaching peak conditions of 550–800 °C and 8–11 kbar. In the high-P/high-T units, granulite and eclogite assemblages developed, with peak conditions ranging between 12 and
Tectonic evolution of the upper allochthon of the Órdenes complex
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Figure 6. Tectonothermal evolution of the allochthonous units of the Órdenes complex, according to the proposed three cycles of convergence and exhumation. (1) Pre-Variscan convergence: D0-M0, formation of a volcanic arc at the northern active margin of Gondwana. D1-M1, accretion of the arc to the northern margin of the Rheic Ocean. D2-M2, tectonic exhumation of the units by extensional detachments. (2) Early Variscan convergence: D3-M3, continued development of the orogenic wedge by understacking of the ophiolitic and basal units, resulting in a new stage of burial and exhumation. (3) Variscan convergence: D4-M4 and D5-M5, representing the final stages of emplacement of the allochthonous complexes, including the gravitational collapse of the orogen. IP—intermediate pressure. Other abbreviations as in Figure 3.
22 kbar and 750–800 °C (Mendia Aranguren, 2000; Arenas and Martínez Catalán, 2002). High-P/high-T fabrics were well preserved in the lowermost units (Sobrado and Belmil units, and also in the Cabo Ortegal complex), whereas in the Fornás and Arinteiro units, D2-M2 retrogression became nearly complete, with D1-M1 association and fabric preserved only in rare, decimeter-scale lenses (Gómez Barreiro, 2004). Differences in the metamorphic gradients experienced by the upper allochthonous units (high-P versus intermediate-P) probably reflect the individualization of two or more tectonic units that underwent different degrees of underthrusting and subduction (Fig. 6). This stage has eluded attempts to establish its age. In the high-P/high-T set, it has been dated in the same units and even with the same samples, at 500–480 Ma (conventional U-Pb isotope dilution and thermal ionization mass spectrometry [IDTIMS]; Peucat et al., 1990; Fernández-Suárez et al., 2002c) and
395–390 Ma (U-Pb SHRIMP; Schäfer et al., 1993; Ordóñez Casado et al., 2001), whereas in the intermediate-P units, monazite ages of 500–490 Ma were interpreted to reflect its age (Abati et al., 1999). If the older ages are those of D1-M1 event, they would imply burial of the upper allochthon closely following or even contemporaneous with widespread gabbro-granodiorite magmatism. Conversely, if the younger ages are correct, they would place D1-M1 temporally far from that Cambro-Ordovician magmatic cycle and in an early Variscan context. Additional information is given by 40Ar/39Ar analyses in hornblendes of D2-M2 amphibolites retrogressive after high-P granulites, yielding ages of ca. 425–410 Ma (Dallmeyer et al., 1997; Gómez Barreiro et al., 2006). If correct, these data would discard the Early Devonian age (395–390 Ma) for the high-P/ high-T event but would not necessarily imply that the CambroOrdovician age is correct.
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A recent secondary ion mass spectrometry (SIMS) U-Pb study of zircons from gabbros, mafic granulites, and migmatitic leucosomes was carried out by Fernández-Suárez et al. (2007), finding that new zircon was growing since at least 410 Ma, in close relation to partial melting. As it is widely accepted that all high-P/high-T units underwent partial melting after having reached their pressure peak (Arenas, 1991; Arenas and Martínez Catalán, 2002), this observation again militates against the Early Devonian age for the D1-M1 stage. The new SIMS data are in accordance with the 40Ar/39Ar analyses and, together, indicate that the peak pressure was reached before 410, and possibly, 425 Ma. D2-M2: Early Exhumation The thickened pile led to gravitational instabilities, resulting in several major extensional detachments in the upper allochthon (Gómez Barreiro, 2004). The best example is the Fornás detachment, marking the contact between the high-P/high-T units and the intermediate-P units in the southwest of the Órdenes complex (Figs. 3–5). A strong regional fabric (S2) developed in the footwall to the detachment (Fornás and Arinteiro units) under amphibolitefacies conditions. The P-T evolution shows isothermal decompression at over 700 °C, which led to partial melting (Gómez Barreiro, 2004) and, subsequently, the roof of the units recorded nearly isobaric cooling at low pressures (<4 kbar). Tectonites developed during decompression are banded mylonites, showing strong shape and crystallographically preferred orientation of amphibole, pyroxene, plagioclase, and quartz, which defines the foliation and a persistent NNW-SSE mineral lineation. Kinematic analysis of the Fornás detachment has revealed a top-tothe-NNW sense of shear (Gómez Barreiro, 2004). Taking into account microstructural features and petrofabric analyses, deformation progressed under conditions bracketed between 750 and 300 °C (Gómez Barreiro, 2004). At the hangingwall to the D2 detachment (O Pino unit), the thermal effects were restricted to the first few meters above the contact, where heat transfer was operative. The rocks experienced heating under low-P and high-T conditions (Fig. 5). Synkinematic leucosomes and mylonites developed at the bottom of the hangingwall, showing a NNW-SSE mineral lineation and kinematic criteria indicating a top-to-the-NNW sense of shear (Gómez Barreiro, 2004). M2 retrogression in the hangingwall to the Fornás detachment completes the counterclockwise loop of the P-T path in the intermediate-P units. In the high-P/high-T units, the geometry of the M1 part of the loop is constrained by P-T conditions deduced from thermobarometry carried out in granulites and eclogites and from corona reactions in granulitic metagabbros of the Sobrado unit (Arenas and Martínez Catalán, 1993, 2002). The M2 part is taken from Díaz García et al. (1999b). In this case, the burial and decompressive trajectories run nearly parallel to each other, not drawing a counterclockwise loop. The first trajectory compares well with the P-T paths modeled for subduction zones by Peacock
(1990, 1991), which are similar, although of opposite sense, to exhumative paths characterized by tectonic denudation followed by erosion, such as M2 in the O Pino unit. Laserprobe 40Ar/39Ar dating of a mylonitic amphibolite from the Fornás detachment constrains its motion around 423 ± 12 Ma (Gómez Barreiro et al., 2006), in accordance with data collected from other units and described in the previous section. D3-M3: A New Cycle of Crustal Shortening and Exhumation The regional fabrics S1 and S2 were overprinted by prograde contractional shear zones and recumbent folds (Figs. 4 and 5), with a vergence and sense of shear to the northeast and east (Martínez Catalán et al., 2002; Gómez Barreiro, 2004; González Cuadra, 2005). In the southwest of the Órdenes complex, typical shear zones led to amphibolite-facies reequilibration with assemblages containing garnet, plagioclase, hornblende, and zoisite. P-T conditions in the highest-grade shear zones were 6–8 kbar and 500–600 °C (Castiñeiras, 2003; Gómez Barreiro, 2004). The previous fabrics were transposed by a new schistosity (S3), axial planar to recumbent folds and show a heterogeneous distribution (Gómez Barreiro, 2004). This event represents a new cycle of thickening in the upper allochthon, but new extensional structures also developed. These are the Corredoiras and Ponte Carreira detachments (Figs. 3 and 5), viewed as the result of gravitational readjustments of the orogenic wedge (Martínez Catalán et al., 1996, 2002; Díaz García et al., 1999b). Together the contractional and subsequent extensional structures define a new, clockwise loop in the P-T paths (Gómez Barreiro, 2004). The timing of the D3-M3 event is constrained by 40Ar/39Ar dating. The amphibolite-facies foliation associated with contractional shear zones has yielded 400–380 Ma (Peucat et al., 1990; Dallmeyer et al., 1991; Gómez Barreiro et al., 2006), whereas the foliation in the extensional detachments of Corredoiras and Ponte Carreira were dated at 375 Ma (Dallmeyer et al., 1997) and 371 Ma (Gómez Barreiro et al., 2006), respectively. These Middle–Late Devonian ages (according to the geological timetable of Gradstein et al., 2004) are coeval with regional amphibolite-facies foliation in the underlying ophiolitic units, where it is also related to thrusting and imbrication (Díaz García et al., 1999a). These contractional structures led to the closure of the oceanic domain between the upper and basal allochthons. Subduction of the latter, thought to represent the outermost edge of the Gondwanan continental margin, has been dated at 385– 375 Ma using 40Ar/39Ar geochronology (Rodríguez et al., 2003). Probably, underthrusting of this continental basement triggered extension of the overlying orogenic wedge, giving rise to the Corredoiras and Ponte Carreira detachments. D4-M4: Emplacement of the Exotic Terranes The exotic terranes were emplaced following pervasive deformation in the autochthon, which gave rise to recumbent folds with axial planar cleavage dated at 359 Ma in the inner parts
Tectonic evolution of the upper allochthon of the Órdenes complex of the belt, close to the allochthonous complexes, and 336 Ma in more external areas (Dallmeyer et al., 1997). The emplacement occurred in more than one step, with the basal allochthon emplaced first along the Lalín-Forcarei thrust above the parautochthon (Fig. 3) at 340 Ma, probably carrying the ophiolites and upper allochthon piggy-back. This movement was followed by renewed thrusting of the two latter at ca. 330–320 Ma (Dallmeyer et al., 1997; Martínez Catalán et al., 2002). The younger thrusts cut across older contractional faults related to the stacking of the different exotic terranes and represent an out-of-sequence thrust system. These thrusts show a topto-the-southeast sense of motion and affected mainly the base of the upper allochthon and the ophiolites. No thermal effects were detected in the upper units, and only retrograde metamorphism occurred in the thrust surface, at very low-T conditions. In the uppermost units, a low-grade crenulation cleavage dated at 330 Ma (Dallmeyer et al., 1997; Gómez Barreiro et al., 2006) is probably related to the activity of the out-of-sequence thrusts. D5-M5: Late Orogenic Collapse Following the D4 contractional event and pervasive shorthening and thickening in the underlying autochthon, gravitational disequilibrium led to renewed gravitational adjustments, giving rise to the Bembibre–Pico Sacro detachment, a complex extensional system that partially reactivated previous out-of-sequence thrusts. Detailed kinematic analysis reveals a shear zone with a general top-to-the-northwest sense of shear. The activity of the Bembibre–Pico Sacro detachment was contemporaneous with extensive melting in the footwall and the establishment of a high-T and lowP gradient (Gómez Barreiro et al., 2002; Gómez Barreiro, 2003), which suggests that the detachment was related to the late gravitational collapse of the orogen (Vanderhaeghe and Teyssier, 2001). The extensive synkinematic magmatism has been dated at 323 ± 11 Ma around the Órdenes complex (Bellido et al., 1992). D6-M6: Late Orogenic Compression Upright folds were formed in conjunction with north-south dextral and sinistral strike-slip shear zones (Fig. 2), and highangle faults overprinted the late extensional detachments and the high-T and low-P isograds. Folds with north-south axial surfaces developed in the Órdenes complex in close relation with a sinistral shear zone developed to the east (Fig. 3). These structures suggest a change in the tectonic regime in which transpression represented an important component of shortening (Iglesias Ponce de León and Choukroune, 1980; Llana-Fúnez and Marcos, 2001). Upright folding has been dated in the autochthon at ca. 314 Ma (Capdevila and Vialette, 1970; Ries, 1979).
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plate and Gondwana indicate that the Iberian authochthon formed part of Gondwana during the whole Paleozoic. The basal allochthon represents the outermost edge of the continent, formed after an episode of continental rifting during the Early Ordovician. That was the time of separation of Avalonia and creation of the Rheic Ocean (Cocks and Fortey, 1988; Scotese and McKerrow, 1990; Winchester et al., 2002) and, for that reason, we suggest that the ocean in front of the Iberian autochthon at that time was also the Rheic or its continuation to the east. The exotic character of the upper allochthon is well established, and its peri-Gondwanan derivation has been discussed above in the geological setting. Drifting occurred at the CambroOrdovician boundary, so that this terrane could be the piece whose separation created the passive margin nowadays preserved in the basal allochthon. However, differences in the zircon populations of the upper allochthon and the autochthon suggest that this supposition might not be true. In any case, both the upper and basal allochthons probably formed part of what has been termed the “Armorica terrane assemblage” (Franke, 2000; Winchester et al., 2002) even when, according to our interpretation, not all the terranes forming the assemblage did actually separate from Gondwana. Orogenic Cycles The tectonometamorphic evolution of the upper allochthon can be readily integrated with what is known from the rest of exotic terranes and the autochthon in a geodynamic model consisting of three orogenic cycles (Fig. 6). Our history starts with the subduction of old oceanic lithosphere existing in front of Gondwana, either that of the Iapetus or Tornquist Ocean, and with pulling apart of the future upper allochthon caused by slab roll-back. Arc magmatism induced the first tectonothermal event identified (D0-M0), dated at the Cambro-Ordovician boundary. Pre-Variscan Cycle The oldest orogenic cycle, which involved subduction, crustal thickening, and exhumation (D1-M1 and D2-M2), resulted presumably from accretion of this peri-Gondwanan terrane to a large continental mass. Its age is poorly constrained between 490 and 410 Ma (Gómez Barreiro et al., 2006; Fernández-Suárez et al., 2007). However, taking into account that in high-P terranes, exhumation closely follows burial, and that the 425–410-Ma U-Pb and 40Ar/39Ar ages reflect exhumation of the high-P/high-T units, a Silurian–Early Devonian age is reasonable for this cycle. In the Early Silurian, Avalonia accreted to Laurentia-Baltica to form Laurussia (Lefort, 1989; Murphy et al., 2004), and this probably was the large continental mass to which the Iberian upper allochthon accreted. This first cycle is considered pre-Variscan because it is related to the closure of either the Iapetus or Tornquist Ocean, and not of the Rheic Ocean.
GEODYNAMIC IMPLICATIONS Sedimentary and faunal evidence, detrital zircon analyses, and the failure to find a suture between the hypothetical Armorica
Early Variscan Cycle The next cycle induced shortening and thickening followed by extension in the upper allochthon (D3-M3), where it started
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after 410 Ma, probably in the Early Devonian. During this cycle, which lasted until the Late Devonian, the oceanic lithosphere of the Rheic Ocean was imbricated and consumed, and the basal allochthon was subducted beneath the orogenic wedge previously formed at the southern margin of Laurussia. This cycle is considered early Variscan because it is related to Laurussia-Gondwana plate convergence and the closure of the Rheic Ocean, but its age (400–370 Ma in the upper allochthon, up to 360 Ma in the ophiolites and basal allochthon; Dallmeyer et al., 1997; Rodríguez et al., 2003) is older than what is normally considered Variscan in a strict sense, that is, essentially Carboniferous. Variscan Cycle The continuation of Laurussia-Gondwana plate convergence once the Rheic Ocean had been closed led to the emplacement of the exotic terranes above the Iberian autochthon (D4-M4). Thrusting, dated between 340 and 320 Ma, took place over the sequences of the passive margin of Gondwana, which had been previously folded and were also shortened by thrust faults. This third cycle started in the Late Devonian, once continental subduction of the basal allochthon was blocked, and lasted during the whole Carboniferous with a fully intracontinental character. It is the only event responsible for the deformation and metamorphism in the parautochthon and autochthon, where it evolved sequentially and diachronously, prograding toward the external zones with time (Dallmeyer et al., 1997). Temperature increases resulting from crustal thickening gave rise to rheological changes in the orogenic wedge and provoked its gravitational collapse (D5-M5). Residual compressional stresses related to oblique plate convergence induced upright folding and transcurrent shearing (D6-M6). Paleogeographic Considerations The model presented in Figure 6 accounts for the orthogonal component of convergence between Laurussia and Gondwana. However, dextral transcurrence between these continental masses during the Devonian and Carboniferous has been widely reported (Gates et al., 1986; Rolet et al., 1994; Van Staal and De Roo, 1995; Franke and Zelazniewicz, 2002; Hatcher, 2002) and has even been considered responsible for the distribution of the different peri-Gondwanan terranes along the European Variscides (Shelley and Bossière, 2000, 2002; Stampfli et al., 2002; Martínez Catalán et al., 2007). A reconstruction of the continental masses at the end of the Variscan cycle is shown in Figure 1, where crosscutting relationships between the Variscan and previous Paleozoic orogenic belts can be appreciated. Northwest Iberia lay adjacent to the Grand Banks of Newfoundland and close to the junction of three Paleozoic collisional belts (the Appalachians, the Scandinavian-British Caledonides, and the German-Polish Caledonides) and of the continental masses involved (Laurentia, Baltica, and Avalonia). As can be seen in Figure 1, the European Variscides lay mostly to the south of Baltica (in present coordinates), and this
reconstruction postdates the strike-slip motion between Laurussia and Gondwana. To remove the effects of dextral transcurrence, most of the Variscan belt should be moved to the upper right corner of Figure 1, thus placing it entirely to the south of Baltica, even using small strike-slip components. This placement implies that the European part of the Variscan belt formed by collision between northern Gondwana, Baltica, and, possibly, eastern Avalonian components of Laurussia. It also suggests that the ophiolitic allochthon of northwest Iberia and equivalent units in central Europe are relicts of the eastern branch of the Rheic Ocean, that the upper allochthon was accreted to southeastern Laurussia during the pre-Variscan orogenic cycle, and that the ocean whose closure allowed the collision was Tornquist. In a general sense, the upper allochthon represents a lateral equivalent of Avalonia and its associated Gander arc, from the point of view of the time of separation from Gondwana and accretion to Laurussia. In both cases, separation took place around the Cambro-Ordovician boundary (Scotese and McKerrow, 1990; Winchester et al., 2002; Martínez Catalán et al., 2007), and accretion to Laurussia occurred in the Early Silurian (Salinian event in Newfoundland and New England; Dunning et al., 1990; Cawood et al., 1995; Hepburn et al., 1995). Figure 7 shows a reconstruction of continents and oceans during the Paleozoic, based on Winchester et al. (2002). The possible location of the European Variscides is outlined, together with that of the exotic terrane presently preserved as the upper allochthon. To constrain these paleopositions, the presence of 1.15-Ga crust in the Eastern Desert of Egypt (Loizenbauer et al., 2001) has been taken into account, as well as the existence of a Mesoproterozoic zircon component in Cambrian sandstone of Israel (Avigad et al., 2003), in Cadomian metagranitoids of the southeast Bohemian Massif (Friedl et al., 2000, 2004), in Early Cambrian sediments of the Małopolska Massif (Belka et al., 2002), and in the Brunovistulian terrane (Hegner and Kröner, 2000). The latter three surround the eastern edge of the Bohemian Massif and are integrated in the Silesian terrane assemblage by Franke and Zelazniewicz (2002), who interpret the involved terranes as derived from Gondwana and accreted to the Baltic margin between the Cambrian and the Silurian. The presence of a Mesoproterozoic zircon population in metasediments of the Iberian autochthon (Fernández-Suárez et al., 2000, 2002a,b; Martínez Catalán et al., 2004) may indicate a paleoposition close to the present northeastern Africa, whereas its absence (Fernández-Suárez et al., 2003) in the upper allochthon would indicate a more western derivation for this exotic terrane. The dextral component of transcurrence between Laurussia and Gondwana might have helped to superimpose the two realms that once were in lateral continuity. CONCLUSIONS The upper allochthon of northwest Iberia represents a periGondwanan terrane drifted away from north Africa around the Cambro-Ordovician boundary and accreted to the Baltic part of
Tectonic evolution of the upper allochthon of the Órdenes complex
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Figure 7. Schematic reconstruction of distribution of continental masses at four stages during the Paleozoic, showing the suggested paleopositions of the European Variscides (dark gray) and the exotic terrane preserved in the upper allochthon (black). Based on Winchester et al. (2002).
Laurussia probably during the Silurian. It registered three orogenic cycles: a pre-Variscan cycle related to its accretion to Laurussia and closure of the intervening Tornquist Ocean, an early Variscan cycle related to Laurussia-Gondwana convergence and closure of the eastern branch of the Rheic Ocean, and a Variscan intracontinental cycle reflecting the Laurussia-Gondwana collision. The pre-Variscan cycle has been identified only in the upper allochthon, whereas the early Variscan cycle was also registered in the ophiolitic and basal allochthons, and the Variscan cycle
affected the whole allochthonous set and is the only cycle developed in the parautochthon and autochthon. Pulling apart from Gondwana of the terrane represented by the upper allochthon was probably induced by slab roll-back of subducting Tornquist oceanic lithosphere, as the terrane registered voluminous arc plutonism, which is at the origin of an early isobaric thermal event. Subsequent burial and exhumation gave rise to a counterclockwise loop in the P-T path in several units of the upper allochthon, which is related to the first, pre-Variscan orogenic
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cycle. A second, clockwise loop is linked to the early Variscan cycle in these units, demonstrating their polyorogenic character. Although the upper allochthon is an exotic terrane drifted away from Gondwana and later emplaced on it, the site of derivation probably differs from that of emplacement. Paleogeographic considerations using available U-Pb zircon ages are compatible with a paleoposition of the northwest Iberian autochthon close to the present northeast Africa and a more western derivation for the upper allochthon. These positions are in agreement with Variscan convergence involving a component of dextral transcurrence. ACKNOWLEDGMENTS This contribution was made possible by funds provided by the Spanish Government agencies Dirección General de Investigación Científica y Técnica and Dirección General de Educacion Superior e Investigación Científica throughout the years and includes data from projects PB97-0234-C02, BTE20010963-C02, and CGL2004-04306-CO2/BTE. JGB acknowledges a postdoctoral fellowship from the Spanish Ministerio de Educación y Ciencia (grant EX2005-0490). The article has benefited from constructive reviews by J. von Raumer and F. Neubauer, and the editorial advice of G. Zulauf. REFERENCES CITED Abalos, B., 1997, Omphacite fabric variation in the Cabo Ortegal eclogite (NW Spain): Relationships with strain symmetry during high-pressure deformation: Journal of Structural Geology, v. 19, p. 621–637, doi: 10.1016/ S0191-8141(97)00001-1. Abalos, B., and Aranguren, A., 1998, Anisotropy of magnetic susceptibility of eclogites: Mineralogical origin and correlation with the tectonic fabric (Cabo Ortegal, Spain): Geodinamica Acta, v. 11, p. 271–283, doi: 10.1016/S0985-3111(99)80017-5. Abalos, B., Azcárraga, J., Gil-Ibarguchi, J.I., Mendía, M., and Santos Zalduegui, J.F., 1996, Flow stress, strain rate and effective viscosity evaluation in a high-pressure metamorphic nappe (Cabo Ortegal, Spain): Journal of Metamorphic Geology, v. 14, p. 227–248. Abati, J., 2002, Petrología metamórfica y geocronología de la unidad culminante del Complejo de Órdenes en la región de Carballo (Galicia, NW del Macizo Ibérico): La Coruña, Spain, Laboratorio Xeolóxico de Laxe, Nova Terra, v. 20, 269 p. Abati, J., Dunning, G.R., Arenas, R., Díaz García, F., González Cuadra, P., Martínez Catalán, J.R., and Andonaegui, P., 1999, Early Ordovician orogenic event in Galicia (NW Spain): Evidence from U-Pb ages in the uppermost unit of the Órdenes complex: Earth and Planetary Science Letters, v. 165, p. 213–228, doi: 10.1016/S0012-821X(98)00268-4. Abati, J., Arenas, R., Martínez Catalán, J.R., and Díaz García, F., 2003, Anticlockwise P-T path of granulites from the Monte Castelo gabbro (Órdenes complex, NW Spain): Journal of Petrology, v. 44, p. 305–327, doi: 10.1093/petrology/44.2.305. Andonaegui, P., González del Tánago, J., Arenas, R., Abati, J., Martínez Catalán, J.R., Peinado, M., and Díaz García, F., 2002, Tectonic setting of the Monte Castelo gabbro (Órdenes complex, northwestern Iberian Massif): Evidence for an arc-related terrane in the hanging wall to the Variscan suture, in Martínez Catalán, J.R., Hatcher, R.D., Arenas, R., and Díaz García, F., eds., Variscan-Appalachian dynamics: The building of the Late Paleozoic basement: Boulder, Colorado, Geological Society of America Special Paper 364, p. 37–56. Araújo, A., and Ribeiro, A., 1997, Estrutura dos domínios meridionais da Zona de Ossa-Morena, in A. Araújo, A., and Pereira, M.F., eds., Estudos de Geologia da Zona de Ossa-Morena (Maciço Ibérico), Livro de homenagem ao Profesor Francisco Gonçalves: Évora, Portugal, p. 169–182.
Arenas, R., 1991, Opposite P, T, t paths of Hercynian metamorphism between the upper units of the Cabo Ortegal complex and their substratum (northwest of the Iberian Massif): Tectonophysics, v. 191, p. 347–364, doi: 10.1016/0040-1951(91)90067-3. Arenas, R., and Martínez Catalán, J.R., 1993, High-pressure and high-temperature metabasites from the Sobrado antiform (northwest of the Iberian Massif, Spain). A petrologic and field study in a granulite-eclogite transition zone: Terra abstracts, Abstract supplement 4 to Terra Nova, v. 5, p. 1. Arenas, R., and Martínez Catalán, J.R., 2002, Prograde development of corona textures in metagabbros of the Sobrado window (Órdenes complex, NW Iberian Massif), in Martínez Catalán, J.R., Hatcher, R.D., Arenas, R., and Díaz García, F., eds., Variscan-Appalachian dynamics: The building of the Late Paleozoic basement: Boulder, Colorado, Geological Society of America Special Paper 364, p. 73–88. Arenas, R., Martínez Catalán, J.R., Sánchez Martínez, S., Díaz García, F., Abati, J., and Fernández-Suárez, J., 2004, Ofiolitas paleozoicas de la sutura varisca de Galicia (NW del Macizo Ibérico, España): distribución, características y significado, in Pereira, E.S., Castroviejo, R., and Ortiz, F., eds., Complejos ofiolíticos en Iberoamérica. Guías de prospección para metales preciosos: Madrid, Ciencia y Tecnología para el Desarrollo, p. 203–230. Arenas, R., Martínez Catalán, J.R., Sánchez Martínez, S., Díaz García, F., Abati, J., Fernández-Suárez, J., Andonaegui, P., and Gómez Barreiro, J., 1997, Paleozoic ophiolites in the Variscan suture of Galicia (northwest Spain): Distribution, characteristics and meaning, in Hatcher, R.D., Jr., Carlson, M.P., McBride, J.H., and Martínez Catalán, J.R., eds., 4-D framework of continental crust: Boulder, Colorado, Geological Society of America Memoir 200, doi: 10.1130/2007.1200(00). Avigad, D., Kolodner, K., McWilliams, M., Persing, H., and Weissbrod, T., 2003, Origin of northern Gondwana Cambrian sandstone revealed by detrital zircon SHRIMP dating: Geology, v. 31, p. 227–230, doi: 10.1130/00917613(2003)031<0227:OONGCS>2.0.CO;2. Belka, Z., Valverde-Vaquero, P., Dörr, W., Ahrendt, H., Wemmer, K., Franke, W., and Schäfer, J., 2002, Accretion of first Gondwana-derived terranes at the margin of Baltica, in Winchester, J.A., Pharaoh, T.C., and Verniers, J., eds., Palaeozoic amalgamation of central Europe: London, Geological Society of London Special Publication 201, p. 19–36. Bellido, F., Brandle, J.L., Lasala, M., and Reyes, J., 1992, Consideraciones petrológicas y cronológicas sobre las rocas graníticas hercínicas de Galicia: Cuadernos do Laboratorio Xeolóxico de Laxe, v. 17, p. 241–261. Caby, R., 1989, Precambrian terranes of Benin-Nigeria and northeast Brazil and the Late Proterozoic south Atlantic fit, in Dallmeyer, R.D., ed., Terranes in the circum-Atlantic Paleozoic orogens: Boulder, Colorado, Geological Society of America Special Paper 230, p. 145–158. Capdevila, R., and Vialette, Y., 1970, Estimation radiométrique de l’âge de la deuxième phase tectonique hercynienne en Galice Moyenne (Nord-Ouest de l’Espagne): Comptes Rendus de l’Académie des Sciences, Paris, v. 270, p. 2527–2530. Castiñeiras, P., 2003, Origen y evolución tectonotermal de las unidades de O Pino y Cariño (Complejos Alóctonos de Galicia) [Ph.D. thesis]: Madrid, Universidad Complutense, 281 p. Cawood, P.A., Van Gool, J.A.M., and Dunning, G.R., 1995, Collisional tectonics along the Laurentian margin of the Newfoundland Appalachians, in Hibbard, J.P., Van Staal, C.R., and Cawood, P.A., eds., Current perspectives in the Appalachian-Caledonian orogen: St. John’s, Newfoundland, Geological Association of Canada Special Paper 41, p. 283–301. Cocks, L.R.M., and Fortey, R.A., 1988, Lower Palaeozoic facies and faunas around Gondwana, in Audley-Charles, M.G., and Hallam, A., eds., Gondwana and the Tethys: London, Geological Society of London Special Publication 37, p. 183–200. Crowley, Q.G., Floyd, P.A., Stedrá, V., Winchester, J.A., Kachlík, V., and Holland, J.G., 2002, The Mariánské-Lázne complex, NW Bohemian Massif: Development and destruction of an early Palaeozoic seaway, in Winchester, J.A., Pharaoh, T.C., and Verniers, J., eds., Palaeozoic amalgamation of central Europe: London, Geological Society of London Special Publication 201, p. 177–195. Dallmeyer, R.D., and Tucker, R.D., 1993, U-Pb zircon age for the Lagoa augen gneiss, Morais complex, Portugal: Tectonic implications: Journal of the Geological Society of London, v. 150, p. 405–410. Dallmeyer, R.D., Ribeiro, A., and Marques, F., 1991, Polyphase Variscan emplacement of exotic terranes (Morais and Bragança massifs) onto Iberian successions: Evidence from 40Ar/39Ar mineral ages: Lithos, v. 27, p. 133–144, doi: 10.1016/0024-4937(91)90025-G.
Tectonic evolution of the upper allochthon of the Órdenes complex Dallmeyer, R.D., Martínez Catalán, J.R., Arenas, R., Gil-Ibarguchi, J.I., Gutiérrez Alonso, G., Farias, P., Aller, J., and Bastida, F., 1997, Diachronous Variscan tectonothermal activity in the NW Iberian Massif: Evidence from 40Ar/39Ar dating of regional fabrics: Tectonophysics, v. 277, p. 307– 337, doi: 10.1016/S0040-1951(97)00035-8. Díaz García, F., 1990, La geología del sector occidental del Complejo de Órdenes (Cordillera Hercínica, NW de España): Coruña, Spain, Laboratorio Xeolóxico de Laxe, Nova Terra, v. 3, 230 p. Díaz García, F., Arenas, R., Martínez Catalán, J.R., González del Tánago, J., and Dunning, G., 1999a, Tectonic evolution of the Careón ophiolite (northwest Spain): A remnant of oceanic lithosphere in the Variscan belt: Journal of Geology, v. 107, p. 587–605, doi: 10.1086/314368. Díaz García, F., Martínez Catalán, J.R., Arenas, R., and González Cuadra, P., 1999b, Structural and kinematic analysis of the Corredoiras detachment: Evidence for early Variscan orogenic extension in the Órdenes complex, NW Spain: International Journal of Earth Sciences, v. 88, p. 337–351, doi: 10.1007/s005310050269. Dunning, G.R., O’Brian, S.J., Colman-Sadd, S.P., Blackwood, R.F., Dickson, W.L., O’Neill, P., and Krogh, T.E., 1990, Silurian orogeny in the Newfoundland Appalachians: Journal of Geology, v. 98, p. 895–913. Farias, P., Gallastegui, G., González-Lodeiro, F., Marquínez, J., Martín Parra, L.M., Martínez Catalán, J.R., de Pablo Maciá, J.G., and Rodríguez Fernández, L.R., 1987, Aportaciones al conocimiento de la litoestratigrafía y estructura de Galicia Central: Memórias da Faculdade de Ciências, Universidade do Porto, v. 1, p. 411–431. Fernández-Suárez, J., Gutiérrez-Alonso, G., Jenner, G.A., and Tubrett, M.N., 1999, Crustal sources in Lower Paleozoic rocks from NW Iberia: Insights from laser ablation U-Pb ages of detrital zircons: Journal of the Geological Society of London, v. 156, p. 1065–1068. Fernández-Suárez, J., Gutiérrez-Alonso, G., Jenner, G.A., and Tubrett, M.N., 2000, New ideas on the Proterozoic–Early Paleozoic evolution of NW Iberia: Insights from U-Pb detrital zircon ages: Precambrian Research, v. 102, p. 185–206, doi: 10.1016/S0301-9268(00)00065-6. Fernández-Suárez, J., Gutiérrez-Alonso, G., Cox, R., and Jenner, G.A., 2002a, Assembly of the Armorica microplate: A strike-slip terrane delivery? Evidence from U-Pb ages of detrital zircons: Journal of Geology, v. 110, p. 619–626, doi: 10.1086/341760. Fernández-Suárez, J., Gutiérrez Alonso, G., and Jeffries, T.E., 2002b, The importance of along-margin terrane transport in northern Gondwana: Insights from detrital zircon parentage in Neoproterozoic rocks from Iberia and Brittany: Earth and Planetary Science Letters, v. 204, p. 75–88, doi: 10.1016/S0012-821X(02)00963-9. Fernández-Suárez, J., Corfu, F., Arenas, R., Marcos, A., Martínez Catalán, J.R., Díaz García, F., Abati, J., and Fernández, F.J., 2002c, U-Pb evidence for a polyorogenic evolution of the HP-HT units of the NW Iberian Massif: Contributions to Mineralogy and Petrolology, v. 143, p. 236–253, doi: 10.1007/s00410-001-0337-2. Fernández-Suárez, J., Díaz García, F., Jeffries, T.E., Arenas, R., and Abati, J., 2003, Constraints on the provenance of the uppermost allochthonous terrane of the NW Iberian Massif: Inferences from detrital zircon U-Pb ages: Terra Nova, v. 15, p. 138–144, doi: 10.1046/j.1365-3121.2003.00479.x. Fernández-Suárez, J., Arenas, R., Abati, J., Martínez Catalán, J.R., Whitehouse, M.J., and Jeffries, T.E., 2007, U-Pb chronometry of polymetamorphic high-pressure granulites: An example from the allochthonous terranes of the NW Iberian Variscan belt, in Hatcher, R.D., Jr., Carlson, M.P., McBride, J.H., and Martínez Catalán, J.R., eds., 4-D framework of continental crust: Boulder, Colorado, Geological Society of America Memoir 200, doi: 10.1130/2007.1200(00). Floor, P., 1966, Petrology of an aegirine-riebeckite gneiss-bearing part of the Hesperian Massif: The Galiñeiro and sourrounding areas, Vigo, Spain: Leidse Geologische Mededelingen, v. 36, p. 1–203. Fonseca, P.E., 1997, Domínios meridionais da Zona de Ossa-Morena e limites com a Zona Sul Portuguesa: metamorfismo de alta pressão relacionado com a sutura Varisca Ibérica, in Araújo, A., and Pereira, M.F., eds., Estudos sobre a Geologia da Zona de Ossa-Morena (Maciço Ibérico), Livro de homenagem ao Profesor Francisco Gonçalves: Évora, Portugal, p. 133–168. Fonseca, P.E., Munhá, J., Pedro, J., Rosas, F., Moita, P., Araújo, A., and Leal, N., 1999, Variscan ophiolites and high-pressure metamorphism in southern Iberia: Ofioliti, v. 24, p. 259–268. Franke, W., 2000, The mid-European segment of the Variscides: Tectonostratigraphic units, terrane boundaries and plate tectonic evolution, in Franke,
329
W., Haak, V., Oncken, O., and Tanner, D., eds., Orogenic processes: Quantification and modelling in the Variscan belt: London, Geological Society of London Special Publication 179, p. 35–61. Franke, W., and Zelazniewicz, A., 2000, The eastern termination of the Variscides: Terrane correlation and kinematic evolution, in Franke, W., Haak, V., Oncken, O., and Tanner, D., eds., Orogenic processes: Quantification and modelling in the Variscan belt: London, Geological Society of London Special Publication 179, p. 63–86. Franke, W., and Zelazniewicz, A., 2002, Structure and evolution of the Bohemian arc, in Winchester, J.A., Pharaoh, T.C., and Verniers, J., eds., Palaeozoic amalgamation of central Europe: London, Geological Society of London Special Publication 201, p. 279–293. Friedl, G., Finger, F., McNaughton, N.J., and Fletcher, I.R., 2000, Deducing the ancestry of terranes: SHRIMP evidence for South America–derived Gondwana fragments in central Europe: Geology, v. 28, p. 1035–1038, doi: 10.1130/0091-7613(2000)028<1035:DTAOTS>2.3.CO;2. Friedl, G., Finger, F., Paquette, J.L., von Quadt, A., McNaughton, N.J., and Fletcher, I.R., 2004, Pre-Variscan geological events in the Austrian part of the Bohemian Massif deduced from U-Pb zircon ages: International Journal of Earth Sciences, v. 93, p. 802–823, doi: 10.1007/s00531-0040420-9. Galán, G., and Marcos, A., 1997, Geochemical evolution of high-pressure mafic granulites from the Bacariza formation (Cabo Ortegal complex, NW Spain): An example of a heterogeneous lower crust: Geologische Rundschau, v. 86, p. 539–555, doi: 10.1007/s005310050162. García Garzón, J., de Pablo Maciá, J.G., and Llamas Borrajo, J.F., 1981, Edades absolutas obtenidas mediante el método Rb/Sr en dos cuerpos de ortoneises en Galicia Occidental: Boletín Geológico y Minero, v. 92, p. 463–466. Gates, A.E., Simpson, C., and Glover, L., III, 1986, Appalachian Carboniferous dextral strike-slip faults: An example from Brookneal, Virginia: Tectonics, v. 5, p. 119–133. Gebauer, D., 1993, Intra-grain zircon dating within the Iberian Massif: Ollo de Sapo augengneisses, bimodal gneisses from the Massif de Guilleries (Girona), graywacke of the Tentudía Group (Serie Negra, SW Spain) and the HP/HT-rock association at Cabo Ortegal (Galicia): Évora, Portugal: Comunicações da XII Reunião de Geologia do Oeste Peninsular, Universidade de Évora, v. 2, p. 41–46. Gee, D.G., and Sturt, B.A., 1985, Scandinavian Caledonides, tectonostratigraphic map, in Gee, D.E., and Sturt, B.A., eds., The Caledonide orogen— Scandinavia and related areas: New York, John Wiley & Sons. Gil-Ibarguchi, J.I., Mendia, M., Girardeau, J., and Peucat, J.J., 1990, Petrology of eclogites and clinopyroxene-garnet metabasites from the Cabo Ortegal complex (northwestern Spain): Lithos, v. 25, p. 133–162, doi: 10.1016/0024-4937(90)90011-O. Girardeau, J., and Gil-Ibarguchi, I., 1991, Pyroxenite-rich peridotites of the Cabo Ortegal complex (northwestern Spain): Evidence for large-scale upper-mantle heterogeneity: Journal of Petrology, v. 32, Special Lherzolites Issue, p. 135–154. Gómez Barreiro, J., 2003, Caracterización cinemática del contacto inferior de la Unidad de Fornás (Complejo de Órdenes, NO del Macizo Ibérico) [Ms. thesis]: Salamanca, Spain, Universidad de Salamanca, 121 pp. Gómez Barreiro, J., 2004, La Unidad de Fornás: Evolución tectonometamórfica del SO del Complejo de Órdenes [Ph.D. thesis]: Salamanca, Spain, Universidad de Salamanca, 294 p. Gómez Barreiro, J., Martínez Catalán, J.R., Arenas, R., and Díaz García, F., 2002, Caracterización cinemática del contacto inferior de la unidad de Fornás (Complejo de Órdenes, NW del Macizo Ibérico): Studia Geologica Salmanticensia, v. 38, p. 105–127. Gómez Barreiro, J., Wijbrans, J.R., Castiñeiras, P., Martínez Catalán, J.R., Arenas, R., Díaz García, F., and Abati, J., 2006, 40Ar/39Ar laserprobe dating of mylonitic fabrics in a polyorogenic terrane of NW Iberia: Journal of the Geological Society of London, v. 163, p. 61–73, doi: 10.1144/0016764905-012. González Cuadra, P., 2005, La Unidad de Corredoiras (Complejo de Órdenes, Galicia): Evolución estructural y metamórfica [Ph.D. thesis]: Salamanca, Spain, Universidad de Salamanca, 254 p. Gradstein, F.M., Ogg, J.G., Smith, A.G., Agterberg, F.P., Bleeker, W., Cooper, R.A., Davydov, V., Gibbard, P., Hinnov, L.A., House, M.R., Lourens, L., Luterbacher, H.P., McArthur, J., Melchin, M.J., Robb, L.J., Shergold, J., Villeneuve, M., Wardlaw, B.R., Ali, J., Brinkhuis, H., Hilgen, F.J., Hooker, J., Howarth, R.J., Knoll, A.H., Laskar, J., Monechi, S., Plumb, K.A., Powell, J., Raffi, I., Röhl, U., Sadler, P., Sanfilippo, A., Schmitz, B.,
330
Gómez Barreiro et al.
Shackleton, N.J., Shields, G.A., Strauss, H., Van Dam, J., van Kolfschoten, T., Veizer, J., and Wilson, D., 2004, A geologic time scale 2004: Cambridge, Cambridge University Press, 589 p. Hargraves, R.B., Dawson, E.M., and Van Houten, F.B., 1987, Palaeomagnetism and age of mid-Palaeozoic ring complexes in Niger, west Africa, and tectonic implications: Geophysical Journal of the Royal Astronomical Society, v. 90, p. 705–729. Hatcher, R.D., Jr., 2002, Alleghanian (Appalachian) orogeny, a product of zipper tectonics: Rotational, transpressive continent–continent collision and closing of ancient oceans along irregular margins, in Martínez Catalán, J.R., Hatcher, R.D., Jr., Arenas, R., and Díaz García, F., eds., VariscanAppalachian dynamics: The building of the Late Paleozoic basement: Boulder, Colorado, Geological Society of America Special Paper 364, p. 199–208. Hegner, E., and Kröner, A., 2000, Review of Nd isotopic data and xenocrystic and detrital zircon ages from the pre-Variscan basement in the eastern Bohemian Massif: Speculations on palinspastic reconstructions, in Franke, W., Haak, V., Oncken, O., and Tanner, D., eds., Orogenic processes, quantification and modelling in the Variscan belt: London, Geological Society of London Special Publication 179, p. 113–130. Hepburn, J.C., Dunning, G.R., and Hon, R., 1995, Geochronology and regional tectonic implications of Silurian deformation in the Nashoba terrane, southeastern New England, U.S.A., in Hibbard, J.P., Van Staal, C.R., and Cawood, P.A., eds., Current perspectives in the Appalachian-Caledonian orogen: St. John’s, Newfoundland, Geological Association of Canada Special Paper 41, p. 349–366. Holdaway, M.J., 1971, Stability of andalusite and the aluminum silicate phase diagram: American Journal of Science, v. 271, p. 97–131. Hubregtse, J.J.M.W., 1973, Petrology of the Mellid area, a Precambrian polymetamorphic rock complex, Galicia, N.W. Spain: Leidse Geologische Mededelingen, v. 49, p. 9–31. Iglesias Ponce de León, M., and Choukroune, P., 1980, Shear zones in the Iberian arc: Journal of Structural Geology, v. 2, p. 63–68, doi: 10.1016/01918141(80)90035-8. Julivert, M., Fontboté, J.M., Ribeiro, A., and Conde, L., 1972, Mapa Tectónico de la Península Ibérica y Baleares, Escala 1: 1.000.000: Madrid, Spain, Instituto Geológico y Minero de España. Kent, D.V., Dia, O., and Sougy, J.M.A., 1984, Paleomagnetism of Lower-Middle Devonian and Upper Proterozoic–Cambrian(?) rocks from Mejeria (Mauritania, west Africa), in Van der Voo, R., Scotese, C.R., and Bonhommet, N., eds., Plate reconstruction from Paleozoic paleomagnetism: Washington, D.C., American Geophysical Union Geodynamics Series 12, p. 99–115. Kuijper, R.P., 1980, Precambrian U-Pb zircon ages from western Galicia (NW Spain): Earth Sciences Review, v. 16, p. 313–316, doi: 10.1016/00128252(80)90051-3. Lancelot, J.R., Allegret, A., and Iglesias-Ponce de León, M., 1985, Outline of Upper Precambrian and Lower Paleozoic evolution of the Iberian Peninsula according to U-Pb dating of zircons: Earth and Planetary Sciences Letters, v. 74, p. 325–337. Leal, N., Pedro, J., Moita, P., Fonseca, P., Araújo, A., and Munhá, J., 1996, Metamorfismo dos sectores meridionais da Zona de Ossa-Morena: actualização de conhecimentos, in A. Araújo, A., and Pereira, M.F., eds., Estudos de Geologia da Zona de Ossa-Morena (Maciço Ibérico), Livro de homenagem ao Profesor Francisco Gonçalves: Évora, Portugal, p. 119–132. Ledru, P., Costa, S., and Echtler, H., 1994, The Massif Central, structure, in Keppie, J.D., ed., Pre-Mezozoic geology in France and related areas, Part III: Berlin, Springer-Verlag, p. 305–323. Lefort, J.P., 1989, Basement correlation across the north Atlantic: Berlin, Springer-Verlag, 148 p. Llana-Fúnez, S., and Marcos, A., 2001, The Malpica-Lamego line: A major crustal-scale shear zone in the Variscan belt of Iberia: Journal of Structural Geology, v. 23, p. 1015–1030, doi: 10.1016/S0191-8141(00)00173-5. Loizenbauer, J., Wallbrecher, E., Fritz, H., Neumayr, P., Khudeir, A.A., and Kloetzli, U., 2001, Structural geology, single zircon ages and fluid inclusion studies of the Meatiq metamorphic core complex: Implications for Neoproterozoic tectonics in the Eastern Desert of Egypt: Precambrian Research, v. 110, p. 357–383, doi: 10.1016/S0301-9268(01)00176-0. Marcos, A., Farias, P., Galán, G., Fernández, F.J., and Llana-Fúnez, S., 2002, Tectonic framework of the Cabo Ortegal complex: A slab of lower crust exhumed in the Variscan orogen (northwestern Iberian Peninsula), in Martínez Catalán, J.R., Hatcher, R.D., Arenas, R., and Díaz García, F.,
eds., Variscan-Appalachian dynamics: The building of the Late Paleozoic basement: Boulder, Colorado, Geological Society of America Special Paper 364, p. 125–142. Martínez Catalán, J.R., Arenas, R., Díaz García, F., Rubio Pascual, F.J., Abati, J., and Marquínez, J., 1996, Variscan exhumation of a subducted Paleozoic continental margin: The basal units of the Órdenes complex, Galicia: NW Spain: Tectonics, v. 15, p. 106–121, doi: 10.1029/95TC02617. Martínez Catalán, J.R., Arenas, R., Díaz García, F., and Abati, J., 1997, Variscan acretionary complex of northwest Iberia: Terrane correlation and succession of tectonothermal events: Geology, v. 25, p. 1103–1106, doi: 10.1130/0091-7613(1997)025<1103:VACONI>2.3.CO;2. Martínez Catalán, J.R., Arenas, R., Díaz García, F., and Abati, J., 1999, Allochthonous units in the Variscan belt of NW Iberia. Terranes and accretionary history, in Sinha, A.K., ed., Basement tectonics 13: Dordrecht, Kluwer, p. 65–84. Martínez Catalán, J.R., Díaz García, F., Arenas, R., Abati, J., Castiñeiras, P., González Cuadra, P., Gómez Barreiro, J., and Rubio Pascual, F., 2002, Thrust and detachment systems in the Órdenes complex (northwestern Spain): Implications for the Variscan-Appalachian geodynamics, in Martínez Catalán, J.R., Hatcher, R.D., Arenas, R., and Díaz García, F., eds., Variscan-Appalachian dynamics: The building of the Late Paleozoic basement: Boulder, Colorado, Geological Society of America Special Paper 364, p. 163–182. Martínez Catalán, J.R., Fernández-Suárez, J., Jenner, G.A., Belousova, E., and Díez Montes, A., 2004, Provenance constraints from detrital zircon U-Pb ages in the NW Iberian Massif: Implications for Paleozoic plate configuration and Variscan evolution: Journal of the Geological Society of London, v. 161, p. 461–473. Martínez Catalán, J.R., Arenas, R., Díaz García, F., Gómez Barreiro, J., Gónzalez Cuadra, P., Abati, J., Castiñeiras, P., Fernández-Suárez, J., Sánchez Martínez, S., Andonaegui, P., Gónzalez Clavijo, E., Díez Montes, A., Rubio Pascual, F.J., and Valle Aguado, B., 2007, Space and time in the tectonic evolution of the northwestern Iberian Massif. Implications for the comprehension of the Variscan belt, in Hatcher, R.D., Jr., Carlson, M.P., McBride, J.H., and Martínez Catalán, J.R., eds., 4-D framework of continental crust: Boulder, Colorado, Geological Society of America Memoir 200, doi: 10.1130/2007.1200(00). Matte, Ph., and Capdevila, R., 1978, Tectonique en grands plis couchés et plissements superposés d’âge hercynien dans la série de Órdenes-Betanzos (Galice Occidentale): Coruña, Spain: Cuadernos del Seminario de Estudios Cerámicos de Sargadelos, v. 27, p. 193–201. Mendia Aranguren, M.S., 2000, Petrología de la Unidad Eclogítica del Complejo de Cabo Ortegal (NW de España): La Coruña, Spain, Laboratorio Xeolóxico de Laxe, Nova Terra, v. 16, 424 p. Murphy, J.B., Fernández-Suárez, J., Keppie, J.D., and Jeffries, T.E., 2004, Contiguous rather than discrete Paleozoic histories for the Avalon and Meguma terranes based on detrital zircon data: Geology, v. 32, p. 585– 588, doi: 10.1130/G20351.1. Neubauer, F., Hoinkes, G., Sassi, F.P., Handler, R., Höck, V., Koller, F., and Frank, W., 1999, Pre-Alpine metamorphism of the eastern Alps: Schweizerische Mineralogische und Petrographische Mitteilungen, v. 79, p. 41–62. Neuman, R.B., and Max, M.D., 1989, Penobscottian-Grampian-Finnmarkian orogenies as indicators of terrane linkages, in Dallmeyer, R.D., ed., Terranes in the circum-Atlantic Paleozoic orogens: Boulder, Colorado, Geological Society of America Special Paper 230, p. 31–45. Ordóñez Casado, B., Gebauer, D., Schäfer, H.J., Gil-Ibarguchi, J.I., and Peucat, J.J., 2001, A single Devonian subduction event for the HP/HT metamorphism of the Cabo Ortegal complex within the Iberian Massif: Tectonophysics, v. 332, p. 359–385, doi: 10.1016/S0040-1951(00)00210-9. Ortega, L.A., Carracedo, M., Larrea, F.J., and Gil-Ibarguchi, J.I., 1996, Geochemistry and tectonic environment of volcanosedimentary rocks from the Ollo de Sapo Formation (Iberian Massif, Spain), in Demaiffe, D., ed., Petrology and geochemistry of magmatic suites of rocks in the continental and oceanic crusts: Brussels, Université Libre de Bruxelles, p. 277–290. Paris, F., and Robardet, M., 1990, Early Paleozoic paleobiogeography of the Variscan regions: Tectonophysics, v. 177, p. 193–213, doi: 10.1016/00401951(90)90281-C. Peacock, S.M., 1990, Numerical simulation of metamorphic pressure-temperature-time paths and fluid production in subducting slabs: Tectonics, v. 9, p. 1197–1211. Peacock, S.M., 1991, Numerical simulation of subduction zone pressure-temperature-time paths: Constraints on fluid production and arc magmatism:
Tectonic evolution of the upper allochthon of the Órdenes complex Philosophical Transactions of the Royal Society, London, v. A335, p. 341–353. Peucat, J.J., Bernard-Griffiths, J., Gil-Ibarguchi, J.I., Dallmeyer, R.D., Menot, R.P., Cornichet, J., and Iglesias Ponce de León, M., 1990, Geochemical and geochronological cross section of the deep Variscan crust: The Cabo Ortegal high-pressure nappe (northwestern Spain): Tectonophysics, v. 177, p. 263–292, doi: 10.1016/0040-1951(90)90285-G. Pin, C., Ortega Cuesta, L.A., and Gil-Ibarguchi, J.I., 1992, Mantle-derived, early Paleozoic A-type metagranitoids from the NW Iberian Massif: Nd isotope and trace-element constraints: Bulletin de la Société Géologique de France, v. 163, p. 483–494. Pin, C., Paquette, J.L., Santos Zalduegui, J.F., and Gil-Ibarguchi, J.I., 2002, Early Devonian supra-subduction zone ophiolite related to incipient collisional processes in the western Variscan belt: The Sierra de Careón unit, Órdenes complex, Galicia, in Martínez Catalán, J.R., Hatcher, R.D., Arenas, R., and Díaz García, F., eds., Variscan-Appalachian dynamics: The building of the Late Paleozoic basement: Boulder, Colorado, Geological Society of America Special Paper 364, p. 57–72. Ploquin, A., and Santallier, D., 1994, The Massif Central, igneous activity, Caledono-Hercynian magmatism in the French Massif Central. Introduction, in Keppie, J.D., ed., Pre-Mezozoic geology in France and related areas, Part III: Berlin, Springer-Verlag, p. 341–343. Rey, P., Burg, J.P., and Casey, M., 1997, The Scandinavian Caledonides and their relationship to the Variscan belt, in Burg, J.P., and Ford, M., eds., Orogeny through time: London, Geological Society of London Special Publication 121, p. 179–200. Ribeiro, A., Pereira, E., and Dias, R., 1990, Central-Iberian zone. Allochthonous sequences. Structure in the northwest of the Iberian Peninsula, in Dallmeyer, R.D., and Martínez García, E., eds., Pre-Mesozoic geology of Iberia: Berlin, Springer-Verlag, p. 220–236. Ries, A.C., 1979, Variscan metamorphism and K-Ar dates in the Variscan fold belt of S Brittany and NW Spain: Journal of the Geological Society of London, v. 136, p. 89–103. Robardet, M., 2002, Alternative approach to the Variscan belt of southwestern Europe: Preorogenic paleobiogeographical constraints, in Martínez Catalán, J.R., Hatcher, R.D., Arenas, R., and Díaz García, F., eds., VariscanAppalachian dynamics: The building of the Late Paleozoic basement: Boulder, Colorado, Geological Society of America Special Paper 364, p. 1–15. Robardet, M., 2003, The Armorica “microplate”: Fact or fiction? Critical review of the concept and contradictory paleobiogeographical data: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 195, p. 125–148, doi: 10.1016/S0031-0182(03)00305-5. Robardet, M., Paris, F., and Racheboeuf, P.R., 1990, Paleogeographic evolution of southwestern Europe during Early Paleozoic times, in McKerrow, W.S., and Scotese, C.R., eds., Paleozoic palaeogeography and biogeography: London, Geological Society of London Memoir 12, p. 411–419. Roberts, D., and Gee, D.G., 1985, An introduction to the structure of the Scandinavian Caledonides, in Gee, D.E., and Sturt, B.A., eds., The Caledonide orogen—Scandinavia and related areas: New York, John Wiley & Sons, p. 55–68. Rocci, G., Bronner, G., and Deschamps, M., 1991, Crystalline basement of the west African craton, in Dallmeyer, R.D., and Lécorché, J.P., eds., The west African orogens and circum-Atlantic correlatives: Berlin, Springer-Verlag, p. 31–61. Rodríguez, J., Cosca, M.A., Gil-Ibarguchi, J.I., and Dallmeyer, R.D., 2003, Strain partitioning and preservation of 40Ar/39Ar ages during Variscan exhumation of a subducted crust (Malpica-Tui complex, NW Spain): Lithos, v. 70, p. 111–139, doi: 10.1016/S0024-4937(03)00095-1. Roger, F., and Matte, Ph., 2005, Early Variscan HP metamorphism in the western Iberian Allochthon–A 390 Ma U-Pb age for the Bragança eclogite (NW Portugal): International Journal of Earth Sciences, v. 94, p. 173–179, doi: 10.1007/s00531-005-0466-3. Rolet, J., Gresselin, F., Jegouzo, P., Ledru, P., and Wyns, R., 1994, Intracontinental Hercynian events in the Armorican Massif, structure and metamorphism, the Variscan orogeny in the Armorican Massif, in Keppie, J.D., ed., Pre-Mesozoic geology in France and related areas: Berlin, SpringerVerlag, p. 195–219. Sánchez Martínez, S., Arenas, R., Andonaegui, P., Martínez Catalán, J.R., and Pearce, J.A., (2007), Geochemistry of two associated ophiolites from the Cabo Ortegal complex (Variscan belt of northwest Spain), in Hatcher, R.D., Jr., Carlson, M.P., McBride, J.H., and Martínez Catalán, J.R., eds.,
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4-D framework of continental crust: Boulder, Colorado, Geological Society of America Memoir 200, doi: 10.1130/2007.1200(00). Santallier, D., Lardeaux, J.M., Marchand, J., and Marignac, Ch., 1994, The Massif Central, metamorphism, in Keppie, J.D., ed., Pre-Mezozoic geology in France and related areas, Part III: Berlin, Springer-Verlag, p. 324–340. Santos, J.F., Schärer, U., Gil-Ibarguchi, J.I., and Girardeau, J., 2002, Genesis of pyroxenite-rich peritotite at Cabo Ortegal (NW Spain): Geochemical and Pb-Sr-Nd isotope data: Journal of Petrology, v. 43, p. 17–43, doi: 10.1093/petrology/43.1.17. Santos Zalduegui, J.F., Schärer, U., and Gil-Ibarguchi, J.I., 1995, Isotope constraints on the age and origin of magmatism and metamorphism in the Malpica-Tuy allochthon, Galicia, NW-Spain: Chemical Geology, v. 121, p. 91–103, doi: 10.1016/0009-2541(94)00123-P. Santos Zalduegui, J.F., Schärer, U., Gil-Ibarguchi, J.I., and Girardeau, J., 1996, Origin and evolution of the Paleozoic Cabo Ortegal ultramafic-mafic complex (NW Spain): U-Pb, Rb-Sr and Pb-Pb isotope data: Chemical Geology, v. 129, p. 281–304, doi: 10.1016/0009-2541(95)00144-1. Schäfer, H.J., Gebauer, D., Gil-Ibarguchi, J.I., and Peucat, J.J., 1993, Ion-microprobe U-Pb zircon dating on the HP/HT Cabo Ortegal complex (Galicia, NW Spain): Preliminary results: Terra Abstracts, v. 5, no. 4, p. 22. Scotese, C.R., 1984, Paleozoic paleomagnetism and the assembly of Pangea, in Van der Voo, R., Scotese, C.R., and Bonhommet, N., eds., Plate reconstruction from Paleozoic paleomagnetism: Washington, D.C., American Geophysical Union Geodynamic Series 12, p. 1–10. Scotese, C.R., and McKerrow, W.S., 1990, Revised world maps and introduction, in McKerrow, W.S., and Scotese, C.R., eds., Paleozoic paleogeography and biogeography: London, Geological Society of London Memoir 12, p. 1–22. Shelley, D., and Bossière, G., 2000, A new model for the Hercynian orogen of Gondwanan France and Iberia: Journal of Structural Geology, v. 22, p. 757–776, doi: 10.1016/S0191-8141(00)00007-9. Shelley, D., and Bossière, G., 2002, Megadisplacements and the Hercynian orogen of Gondwanan France and Iberia, in Martínez Catalán, J.R., Hatcher, R.D., Arenas, R., and Díaz García, F., eds., Variscan-Appalachian dynamics: The building of the Late Paleozoic basement: Boulder, Colorado, Geological Society of America Special Paper 364, p. 209–222. Stampfli, G.M., and Borel, G.D., 2002, A plate tectonic model for the Paleozoic and Mesozoic constrained by dynamic plate boundaries and restored synthetic oceanic isochrons: Earth and Planetary Sciences Letters, v. 196, p. 17–33, doi: 10.1016/S0012-821X(01)00588-X. Stampfli, G.M., von Raumer, J.F., and Borel, G.D., 2002, Paleozoic evolution of pre-Variscan terranes: From Gondwana to the Variscan collision, in Martínez Catalán, J.R., Hatcher, R.D., Arenas, R., and Díaz García, F., eds., Variscan-Appalachian dynamics: The building of the Late Paleozoic basement: Boulder, Colorado, Geological Society of America Special Paper 364, p. 263–280. Stephens, B., and Gee, D.G., 1989, Terranes and polyphase accretionary history in the Scandinavian Caledonides, in Dallmeyer, R.D. ed., Terranes in the circum-Atlantic Paleozoic orogens: Boulder, Colorado, Geological Society of America Special Paper 230, p. 17–30. Valverde Vaquero, P., and Dunning, G.R., 2000, New U-Pb ages for Early Ordovician magmatism in central Spain: Journal of the Geological Society of London, v. 157, p. 15–26. Van Calsteren, P.W.C., Boelrijk, N.A.I.M., Hebeda, E.H., Priem, H.N.A., Den Tex, E., Verdurmen, E.A.T.H., and Verschure, R.H., 1979, Isotopic dating of older elements (including the Cabo Ortegal mafic-ultramafic complex) in the Hercynian orogen of NW Spain: Manifestations of a presumed Early Paleozoic mantle-plume: Chemical Geology, v. 24, p. 35–56, doi: 10.1016/0009-2541(79)90011-1. Vanderhaeghe, O., and Teyssier, C., 2001, Crustal-scale rheological transitions during late-orogenic collapse: Tectonophysics, v. 335, p. 211–228, doi: 10.1016/S0040-1951(01)00053-1. Van der Voo, R., 1982, Pre-Mesozoic paleomagnetism and plate tectonics: Annual Reviews of Earth and Planetary Science, v. 10, p. 191–220. Van der Voo, R., 1988, Paleozoic paleogeography of North America, Gondwana, and intervening displaced terranes: Comparisons of paleomagnetism with paleoclimatology and biogeographical patterns: Geological Society of America Bulletin, v. 100, p. 311–324, doi: 10.1130/00167606(1988)100<0311:PPONAG>2.3.CO;2. Van Staal, C.R., and De Roo, J.A., 1995, Mid-Paleozoic tectonic evolution of the Appalachian central mobile belt in northern New Brunswick, Canada: Collision, extensional collapse and dextral transpression, in Hibbard,
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Gómez Barreiro et al.
J.P., Van Staal, C.R., and Cawood, P.A., eds., Current perspectives in the Appalachian-Caledonian orogen: St. John’s, Newfoundland, Geological Association of Canada Special Paper 41, p. 367–389. Van Zuuren, A., 1969, Structural petrology of an area near Santiago de Compostela (NW Spain): Leidse Geologische Mededelingen, v. 45, p. 1–71. Vialette, Y., Casquet, C., Fúster, J.M., Ibarrola, E., Navidad, M., Peinado, M., and Villaseca, C., 1987, Geochronological study of orthogneisses from the Sierra de Guadarrama (Spanish central system): Neues Jahrbuch für Mineralogie Monatshefte, v. 10, p. 465–479. Vogel, D.E., 1967, Petrology of an eclogite- and pyrigarnite-bearing polymetamorphic rock complex at Cabo Ortegal: NW Spain: Leidse Geologische Mededelingen, v. 40, p. 121–213.
von Raumer, J.F., Stampfli, G.M., and Bussy, F., 2003, Gondwana-derived microcontinents—The constituents of the Variscan and Alpine collisional orogens: Tectonophysics, v. 365, p. 7–22, doi: 10.1016/S00401951(03)00015-5. Winchester, J.A., Pharaoh, T.C., and Verniers, J., 2002, Palaeozoic amalgamation of central Europe: An introduction and synthesis of new results from recent geological and geophysical investigations, in Winchester, J.A., Pharaoh, T.C., and Verniers, J., eds., Palaeozoic amalgamation of central Europe: London, Geological Society of London Special Publication 201, p. 1–18. MANUSCRIPT ACCEPTED BY THE SOCIETY 3 OCTOBER 2006
Printed in the USA
Geological Society of America Special Paper 423 2007
Crustal growth and deformational processes in the northern Gondwana margin: Constraints from the Évora Massif (Ossa-Morena zone, southwest Iberia, Portugal) M. Francisco Pereira* Centro de Geofísica de Évora, Departamento de Geociências, Universidade de Évora, 7002-554 Évora, Portugal J. Brandão Silva Departamento de Geologia, Faculdade de Ciências, Universidade de Lisboa, 1749-016 Lisbon, Portugal Martim Chichorro Patrícia Moita Centro de Geofísica de Évora, Departamento de Geociências, Universidade de Évora, 7002-554 Évora, Portugal José F. Santos Departamento de Geociências, Universidade de Aveiro, 3810-193 Aveiro, Portugal Arturo Apraiz Departamento de Geodinámica, Facultad de Ciencias y Tecnología, Euskal Herriko Universidad del Pais Vasco, Bilbao, Spain Cristina Ribeiro Departamento de Geociências, Universidade de Évora, 7002-554 Évora, Portugal
ABSTRACT The aim of this article is to present a compilation of available information on the Évora Massif based on structural mapping, whole-rock geochemistry, recognition of metamorphic mineral assemblages, and geothermobarometry. In our view, transcurrent movements responsible for strong orogen-parallel stretching were dominant and had a major role in the geodynamic evolution of this part of Ossa-Morena zone (southwest Iberian Massif). Cadomian and Variscan orogenic events separated by a period of intense rifting were the cause for the composite distribution of zones with contrasting metamorphic paths, the structural complexity, the variety of lithological associations, and the sequence of deformation events and magmatism. The proposed geodynamic reconstruction for this segment of the northern Gondwana continental margin includes three main stages in chronological order: (1) Neoproterozoic accretion and continental magmatic arc developing, dismantling, and reworking, followed
*E-mail:
[email protected]. Pereira, M.F., Silva, J.B., Chichorro, M., Moita, P., Santos, J.F., Apraiz, A., and Ribeiro, C., 2007, Crustal growth and deformational processes in the northern Gondwana margin: Constraints from the Évora Massif (Ossa-Morena zone, southwest Iberia, Portugal), 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. 333–358, doi: 10.1130/2007.2423(16). For permission to copy, contact
[email protected]. ©2007 Geological Society of America. All rights reserved.
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Pereira et al. by late-“orogenic” magmatism; (2) Lower Paleozoic crustal thinning, block tilting, and mantle upwelling, induced by generalized rifting, leading to the formation of marine basins with carbonate platform sediments and thick accumulations of volcaniclastic and terrigenous sediments, contemporaneous with normal and enriched mid-oceanic ridge basalt–type magmatism; and (3) Upper Paleozoic transpressional orogenesis resulting from obliquity of convergence and the geometry of the involved blocks. The third stage includes the tectonic inversion of Lower Paleozoic basins, crustal thickening, the exhumation of high- to medium-pressure rocks and partial exhumation of high-grade metamorphic lithologies (controlled by local transtension and major detachments), the formation of synorogenic basins filled with volcanicsedimentary sequences, and finally, the emplacement of late Variscan granodiorites and granites. Keywords: Early Paleozoic rifting, Variscan orogeny, transcurrent tectonics, highgrade terranes, local transtension
INTRODUCTION Although not totally understood, the advent of Paleozoic rifting that culminated in the opening of the Rheic Ocean (named after the Greek mythical goddess Rhea) and correlative basins began in Early Cambrian time with the breaking up of a Neoproterozoic continental margin (e.g., Moores and Twiss, 1995; Windley, 1997). Meanwhile, in Europe a preserved Neoproterozoic continental magmatic arc (Cadomian orogeny) is clearly indicated, with no evidence found, so far, for continental collision with its associated high-pressure metamorphism. After this accretion, the northern Gondwana margin began to fragment in the Early Cambrian, giving rise to the development of marine basins with important and widespread magmatism (Floyd et al., 2000; Dörr et al., 2002; Linnemann and Romer, 2002; Murphy et al., 2002; Linnemann et al., 2004; Silva and Pereira, 2004). It is also firmly established that an Upper Paleozoic accretionsubduction process was responsible for the closure and tectonic inversion of these marine/oceanic basins, caused by the collision of Gondwana and Laurussia (Variscan orogeny; Ribeiro et al., 1990; Matte, 1991, 2001; Paris and Robardet, 1994; Dias and Ribeiro, 1995; Tait et al., 1997; Castro et al., 1999; Díaz Garcia et al., 1999; Franke et al., 2000, and references therein; Shelley and Bossière, 2000; Stampfli, 2000; Martínez Catalán et al., 2002, and references therein). Despite this accepted geodynamic scheme there is no real consensus about the mode and timing of rifting, oceanization, kinematics, subduction, thickening, and collapse-related processes that occurred in this margin of Gondwana during a period of more than 250 m.y. of Earth’s history (see discussions in Pereira and Silva, 2001; Shelley and Bossière, 2001, 2002; Cartier et al., 2002; Robardet, 2002, 2003). Until now the available information about the extensional structures related to crustal thinning, the distribution of anorogenic magmatism and hypothetical oceanic crust remnants, the plausible existence of more than one suture zone and associated high-pressure rocks, the occurrence of orogenic continental arc-related magmatic complexes, and the reconstruction of subduction-related
thrust tectonics and high-pressure units was insufficient to create a well-established geodynamic model. Our study describes the main geological features of the Évora Massif. This unit records a complex geological evolution in which widespread high-grade metamorphism and plutonism prevailed, and major ductile transcurrent shear zones controlled the overall structure. The detailed field data, mesoscopic- and microscopic-scale structural information, geochemical characterization of magmatism (based on immobile elements), and metamorphic petrology (based on mineral chemistry presented here) clearly show that this Portuguese domain of the Ossa-Morena zone (part of the Iberian Massif) is important to understand the overall geodynamic evolution of the northern Gondwana continental margin. New data presented here were interpreted in the light of the concept of transpression and transtension (e.g., Dewey et al., 1998) to explain a complex Paleozoic history. This history includes initial continued fragmentation of the Cadomian basement and creation of a network of seaways with mid-oceanic ridge basalt (MORB)-like composition magmatism followed by basin inversions, accretion, subduction with orogenic magmatism, extensional collapse of the chain, and basin development. GEOLOGICAL SETTING, RELATION BETWEEN DEFORMATION AND METAMORPHISM The Évora Massif, first defined by Carvalhosa (1983), and later referred by Quesada and Munhá (1990), Gonçalves and Carvalhosa (1994), and Pereira et al. (2003a), represents an important domain of the western Ossa-Morena zone where the metamorphic grade varies from medium to high, including a 25- to 45-km-wide and 75-km-long complex crystalline core zone. The different rock types from this domain are affected by strain partitioning, with the presence of important transcurrent zones separating different tectonic units (Fig. 1): (1) the Évora high-grade metamorphic terrains, (2) the Montemor-o-Novo shear zone, and (3) the Évora medium-grade metamorphic terrains.
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Granites and granodiorites (Carboniferous) Felsic porphyres (Carboniferous) Late brittle/semi-ductile shear faults Ductile transcurrent shear faults Ductile normal faults
Figure 1. Simplified geological map of the Évora Massif showing the main tectonic units: Montemor-o-Novo shear zone, Évora high-grade metamorphic terrains, and Évora medium-grade metamorphic terrains. Letters A–J mark locations of cross-sections that are shown in Figure 3. Modifed from Carvalhosa et al. (1969, 1987), Oliveira (1992), Carvalhosa and Zbyszewski (1994), Carvalhosa (1999), Pereira and Silva (2002), and Pereira et al. (2003a).
In this study we mainly used the characteristics of metamorphosed basic rocks to obtain preliminary data from P-T conditions by conventional thermobarometric calculations. Analyses of mineral chemistry were performed at the University of Oviedo, Spain, using a CAMEBAX SX50 microprobe, and at the Oporto Laboratory of the Instituto de Energia, Tecnologia e Inovação, Portugal (Tables 1–4).
The Évora Massif was mainly affected by a clockwise metamorphic path with a distinct series of metamorphic facies indicative of a complex and diverse distribution of tectonic processes involving crustal extension under a transcurrent regime of deformation. Previous structures and high-pressure metamorphism that allegedly would exist related to crustal thickening have been almost completely erased.
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TABLE 1A. MAJOR AND TRACE ELEMENT CONTENTS OF THE ÉVORA MASSIF MAGMATIC ROCKS: BASIC ROCKS (GROUPS I AND II, AMPHIBOLITES) Sample CSN-A24 BSC-2 BSC-3 VAL-2 EVR-4 BSC-10 SCD ARL-1 ARL-2 ARL-3 ARL-10 BDV-1 BDV-2/R (wt%) SiO2 51.81 51.93 51.74 48.22 47.57 49.33 47.19 47.82 48.76 47.71 48.5 44.2 49.11 TiO2 2.454 3.79 3.13 1.852 1.7 1.62 1.884 2.059 1.533 1.37 1.81 1.389 1.944 Al2O3 13.58 14.13 12.73 15.4 15.21 14.58 15.45 14.64 14.78 15.14 14.8 14.1 14.48 Fe2O3 15.41 13.74 17.02 12.78 12.09 10.88 11.59 12.16 9.91 10.66 11.24 11.42 9.81 MnO 0.262 0.25 0.344 0.191 0.18 0.184 0.187 0.212 0.164 0.159 0.173 0.16 0.164 MgO 5.13 3.78 3.34 5.29 7.89 6.83 6.93 7.43 7.74 8.05 7.17 12.91 6.62 CaO 6.98 7.09 6.37 10.55 10.16 9.87 12.51 10.65 11.43 10.73 10.63 9.99 13.05 Na2O 3.02 4.04 3.59 3.67 2.8 2.9 2.6 2.92 3.01 3.08 2.82 1.94 3.29 K2O 0.11 0.23 0.55 0.66 0.15 0.73 0.72 0.31 0.28 0.51 0.45 0.32 0.09 P2O5 0.23 0.42 0.95 0.27 0.13 0.2 0.2 0.24 0.18 0.15 0.22 0.13 0.29 Total 99.47 100.0 100.0 99.69 98.7 98.68 99.92 99.59 98.84 98.89 98.93 98.79 99.4 (ppm) Rb 3.0 4.0 14.0 11.0 <2.0 22.0 8.0 4.0 4.0 10.0 9.0 13.0 2.0 Cs 0.5 <0.5 0.6 0.9 <0.5 <0.5 0.2 0.4 0.4 1.0 1.0 1.0 0.9 Be 1.0 2.0 2.0 <1.0 1.0 1.0 1.0 1.0 <1.0 1.0 1.0 2.0 2.0 Sr 74.0 162.0 100.0 181.0 157.0 194.0 229.0 96.0 200.0 157.0 96.0 187.0 394.0 Ba 186.0 176.0 122.0 175.0 20.0 643.0 112.0 51.0 39.0 122.0 85.0 50.0 61.0 Sc 46.0 46.0 41.0 43.0 44.0 43.0 39.0 42.0 40.0 40.0 40.0 37.0 36.0 V 405.0 443.0 311.0 281.0 292.0 264.0 270.0 328.0 274.0 270.0 350.0 253.0 289.0 Cr 39.0 33.0 <20.0 253.0 254.0 301.0 121.0 242.0 330.0 420.0 202.0 1020.0 330.0 Co 37.0 28.0 27.0 47.0 37.0 32.0 51.0 42.0 39.0 43.0 31.0 58.0 37.0 Ni 13.0 <20.0 <20.0 70.0 71.0 30.0 77.0 90.0 106.0 147.0 48.0 402.0 128.0 Ga 22.0 22.0 24.0 19.0 17.0 17.0 22.0 17.0 15.0 15.0 18.0 17.0 20.0 Y 46.8 67.0 81.0 33.9 31.0 32.0 27.6 33.7 23.4 22.0 29.0 24.9 30.4 Nb 3.4 11.0 13.0 3.9 3.0 4.0 9.8 8.6 9.6 6.0 8.0 7.1 20.8 Ta 0.19 0.8 0.8 0.1 0.1 0.2 0.63 0.61 0.68 0.4 0.5 0.43 1.48 Zr 150.0 311.0 366.0 94.0 115.0 134.0 115.0 131.0 93.0 83.0 115.0 88.0 156.0 Hf 4.6 8.3 9.0 3.1 3.2 3.5 3.1 3.7 2.7 2.4 3.2 2.5 4.1 U 0.32 1.1 1.2 0.14 0.4 0.4 0.24 0.34 0.29 0.2 0.3 0.17 0.5 Th 0.77 2.6 3.2 0.28 0.5 0.8 0.69 0.79 0.8 0.5 1.0 0.48 1.75 La 8.11 9.8 29.0 3.99 4.5 8.0 11.6 7.75 8.61 5.6 9.0 5.77 17.9 Ce 19.6 45.1 66.7 11.1 12.3 19.2 27.6 18.5 17.8 13.8 21.3 12.9 33.9 Pr 3.19 6.52 9.59 2.0 2.07 2.9 3.56 2.98 2.63 2.03 3.1 2.04 4.63 Nd 17.7 32.8 47.5 11.3 11.6 14.8 17.5 15.8 12.8 10.4 15.0 10.8 21.3 Sm 5.93 9.8 13.0 4.15 3.8 4.4 4.84 4.96 3.75 3.1 4.4 3.51 5.61 Eu 2.08 2.98 3.91 1.63 1.5 1.48 1.74 1.68 1.32 1.16 1.63 1.32 1.87 Gd 6.93 11.3 14.5 5.04 5.1 5.5 5.23 5.36 3.94 3.9 5.4 3.86 5.58 Tb 1.31 2.1 2.6 0.96 0.9 1.0 0.86 0.97 0.7 0.7 0.9 0.71 0.95 Dy 8.63 12.8 15.4 6.73 5.9 6.0 5.11 6.28 4.33 4.5 5.6 4.54 5.94 Ho 1.83 2.6 3.1 1.45 1.2 1.2 1.0 1.29 0.9 0.9 1.2 0.95 1.19 Er 5.47 7.8 9.0 4.35 3.7 3.7 2.73 3.82 2.67 2.8 3.4 2.78 3.44 Tm 0.801 1.15 1.33 0.639 0.55 0.54 0.38 0.558 0.389 0.41 0.5 0.401 0.499 Yb 5.11 7.2 8.2 3.88 3.5 3.3 2.39 3.56 2.52 2.6 3.1 2.59 3.23 Lu 0.802 1.01 1.13 0.571 0.49 0.48 0.368 0.55 0.385 0.37 0.44 0.399 0.499
The main pervasive foliation strikes N310° but is locally rotated to north-south by the effect of ductile deflections and subsequently brittle deformation. This planar fabric is habitually associated to a strong mylonitization and affected by folding. Locally it is overprinted by a later subvertical foliation. Continuing sinistral shearing acting on distinct structural levels created a conspicuous structural pattern, with fold axes subparallel to a subhorizontal- to moderately dipping stretching lineation that strikes from N330°–340° to N300°–310°, along which the main tectonic transport took place (Chichorro, 2006). The heterogeneous distribution of deformation caused the individualization of different
REL 46.53 1.666 15.55 10.67 0.165 6.51 13.59 2.28 0.32 0.21 98.94 6.0 0.5 <1.0 215.0 52.0 42.0 320.0 209.0 35.0 64.0 20.0 27.8 6.7 0.49 108.0 3.1 0.29 0.66 7.4 16.0 2.48 12.5 3.9 1.42 4.32 0.8 5.23 1.07 3.22 0.471 2.98 0.464
sectors bounded by transcurrent faults, showing distinct degrees of strain and structure development (Pereira et al., 2003a). Montemor-o-Novo Shear Zone The Montemor-o-Novo shear zone, which extends along ~30 km from Cabrela to Boa Fé (see Fig. 1 for location) and is 2–10 km wide, is made of strongly sheared rocks under amphibolitic-facies metamorphic conditions (Chichorro et al., 2003, 2004). This ductile shear zone (Figs. 2 and 3) includes Ediacaran black metacherts, metapelites, metagraywackes, mica schists,
Crustal growth and deformational processes in the northern Gondwana margin
Sample (wt%) SiO2 TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5 Total (ppm) Rb Cs Be Sr Ba Sc V Cr Co Ni Ga Y Nb Ta Zr Hf U Th La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
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TABLE 1B. MAJOR AND TRACE ELEMENT CONTENTS OF THE ÉVORA MASSIF MAGMATIC ROCKS: BASIC ROCKS (GROUP III, AMPHIBOLITES) AND FELSIC ROCKS (GRANITOIDS) MAR SLV MOF BSC-6 CTO VAL-1 CSN-B24 PFPAS PFAA1 GPAS PFMAS 51.81 2.18 12.8 13.58 0.158 4.5 5.37 4.11 0.01 0.19 99.31
51.01 2.179 15.19 11.2 0.208 5.57 4.93 5.03 1.38 0.21 99.37
47.33 2.112 15.13 11.31 0.19 7.87 9.33 3.68 0.79 0.18 99.68
<1.0 0.1 1.0 80.0 131.0 46.0 288.0 154.0 42.0 52.0 11.0 35.1 4.1 0.22 123.0 3.6 0.23 0.26 5.63 14.5 2.4 13.1 4.53 1.45 5.34 1.03 7.02 1.52 4.49 0.667 4.06 0.597
16.0 0.3 1.0 73.0 224.0 46.0 294.0 99.0 48.0 38.0 16.0 34.5 4.7 0.3 138.0 4.1 0.15 0.31 6.1 16.7 2.82 15.0 5.05 1.64 5.58 1.03 6.72 1.42 4.15 0.589 3.71 0.575
15.0 0.7 1.0 296.0 134.0 41.0 292.0 275.0 33.0 87.0 19.0 37.4 4.4 0.22 127.0 3.8 0.2 0.28 5.89 15.0 2.59 14.6 5.01 1.9 5.83 1.08 7.05 1.5 4.33 0.61 3.88 0.593
48.52 2.02 14.62 12.55 0.206 7.21 9.33 3.3 0.37 0.18 99.27 9.0 <0.5 1.0 160.0 89.0 50.0 307.0 259.0 45.0 31.0 2.0 38.0 4.0 0.2 126.0 3.6 0.2 0.2 6.8 17.5 2.65 14.6 4.7 1.67 6.0 1.1 7.0 1.4 4.3 0.64 4.0 0.56
48.38 1.436 13.6 10.55 0.179 6.77 14.46 2.1 0.65 0.12 99.2
49.42 0.979 13.57 9.09 0.151 10.19 14.27 1.51 0.13 0.08 100.4
49.05 2.193 13.7 14.02 0.227 6.44 8.06 2.33 1.91 0.22 99.58
23.0 0.7 <1.0 249.0 146.0 38.0 279.0 407.0 44.0 168.0 15.0 27.7 4.0 0.16 78.0 2.4 0.19 0.17 3.63 8.93 1.59 9.06 3.26 1.26 3.94 0.75 5.16 1.14 3.4 0.498 3.04 0.446
5.0 0.5 <1.0 147.0 25.0 55.0 260.0 1430.0 44.0 118.0 16.0 20.7 4.6 0.1 51.0 1.9 0.07 0.12 2.18 5.96 0.94 5.66 2.07 0.808 2.69 0.5 3.12 0.67 2.03 0.304 1.92 0.299
72.0 0.8 1.0 129.0 359.0 45.0 344.0 63.0 42.0 30.0 20.0 36.9 4.0 0.28 124.0 3.7 0.23 0.55 8.45 18.9 2.93 15.9 5.06 1.87 5.87 1.08 6.94 1.43 4.3 0.633 4.01 0.613
paragneisses, amphibolites, and felsic gneisses (Serie Negra or Escoural Formation); a Lower-Middle Cambrian igneous (felsic-dominated)-sedimentary complex with marbles, interbedded felsic and mafic metavolcanics, and mica schists (Monfurado Formation); and an Middle–Upper Cambrian–Lower Ordovician(?) igneous (basic-dominated)-sedimentary complex made of amphibolites with mica schists, metatuffs, and calc-silicate rocks (Carvalhal Formation) (e.g., Carvalhosa and Zbyszewski, 1994). Meter-scale lenses of eclogites occur, associated with mediumpressure felsic gneisses and amphibolites (Pedro, 1996; Leal et al., 1997; Leal, 2001).
70.6 0.355 14.82 2.59 0.056 0.83 2.21 3.58 3.87 0.15 99.65 177.0 12.4 4.0 197.0 692.0 6.0 25.0 <20.0 4.0 <20.0 15.0 14.0 10.0 1.1 143.0 4.1 2.7 23.0 29.6 56.0 6.0 21.1 4.2 0.89 3.3 0.5 2.7 0.5 1.5 0.22 1.4 0.2
67.82 0.463 15.36 3.37 0.059 1.12 2.99 3.96 3.01 0.18 98.98 131.0 8.8 4.0 279.0 918.0 7.0 43.0 <20.0 5.0 <20.0 16.0 10.0 9.0 0.6 197.0 4.9 2.6 18.3 36.2 63.5 6.39 21.7 3.6 1.05 2.7 0.4 2.1 0.4 1.1 0.16 1.0 0.15
75.99 0.03 14.04 0.77 0.054 0.08 0.63 3.9 3.9 0.18 100.19
71.98 0.327 14.49 2.46 0.054 0.77 2.1 3.54 3.78 0.13 100.14
287.0 28.1 16.0 18.0 16.0 2.0 <5.0 <20.0 <1.0 <20.0 14.0 9.0 140 3.9 21.0 1.3 9.7 1.3 2.9 6.4 0.73 3.6 1.2 0.12 1.4 0.3 1.6 0.3 0.7 0.1 0.6 0.07
156.0 9.1 4.0 156.0 483.0 6.0 32.0 <20.0 3.0 <20.0 15.0 15.0 9.0 1.0 140.0 4.0 2.9 21.3 27.2 51.3 5.58 19.8 4.0 0.82 3.4 0.5 2.8 0.5 1.7 0.26 1.6 0.24
Overlying this strongly deformed and metamorphosed basement, the most significant lithologies are Early Carboniferous detrital sediments and volcanics (pelites, graywackes, conglomerates with associated andesitic to trachyandesitic tuffs, dacitic and andesitic flows, and polygenic conglomerates; Cabrela Formation; Ribeiro, 1983; Oliveira et al., 1991; Carvalhosa and Zbyszewski, 1994; Pereira and Oliveira, 2001). This Viséan detrital sedimentation includes olistostromes of Middle Devonian limestones and calciturbidites (Pereira and Oliveira, 2003). Here the deformation is represented by an incipient slaty cleavage or locally, in the southern boundary with the basement, by a
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TABLE 1C. MAJOR AND TRACE ELEMENT CONTENTS OF THE ÉVORA MASSIF: SEDIMENTARY ROCKS Sample MTN-1 REL-10 XC-20 BDV VAL-23 SN (wt%) SiO2 74.2 57.9 72.3 48.0 60.4 61.605 TiO2 0.72 0.80 0.64 1.196 0.81 0.7935 Al2O3 13.5 20.0 14.2 28.0 17.7 17.305 Fe2O3 3.90 8.17 4.44 7.74 5.86 6.54 MnO 0.03 0.20 0.04 0.05 0.06 0.063 MgO 0.71 1.63 0.97 2.19 2.58 2.8875 CaO 0.03 0.03 0.22 0.24 1.88 1.1625 Na2O 1.32 1.07 1.03 0.48 4.62 3.5475 K2O 3.26 4.09 3.59 5.97 2.24 3.51 P2O5 0.04 0.05 0.09 0.11 0.14 0.21 Total 100.1 99.04 99.58 99.41 98.92 100.2 (ppm) Rb 129.0 178.0 141.0 246.0 73.0 124.75 Cs N.D. N.D. N.D. N.D. N.D. N.D. Be N.D. N.D. N.D. N.D. N.D. N.D. Sr 30.0 87.0 80.0 134.0 133.0 145.0 Ba 608.0 872.0 710.0 924.0 769.0 905.5 Sc 8.0 18.0 9.0 22.0 18.0 17.75 V N.D. N.D. N.D. N.D. N.D. N.D. Cr N.D. N.D. N.D. N.D. N.D. N.D. Co N.D. N.D. N.D. N.D. N.D. N.D. Ni 0.0 30.0 0.0 30.0 40.0 45.25 Ga 15.0 23.0 16.0 33.0 20.0 22.25 Y 34.0 45.0 30.0 47.0 35.0 29.75 Nb 15.0 15.0 13.0 23.0 12.0 15.75 Ta N.D. N.D. N.D. N.D. N.D. N.D. Zr 508.0 121.0 281.0 190.0 171.0 180.75 Hf 13.00 4.00 7.00 5.00 5.00 5.025 U 4.0 2.0 3.0 3.0 3.0 3.525 Th 20.0 16.0 12.0 22.0 12.0 10.775 La 39.2 69.1 33.7 74.2 39.3 41.85 Ce 80.8 128.0 71.1 148.0 77.5 62.455 Pr 8.67 15.50 7.51 16.40 9.01 8.875 Nd 31.4 65.5 28.2 60.8 35.3 33.9 Sm 5.6 10.8 5.1 10.8 6.7 6.825 Eu 1.2 2.2 1.2 2.1 1.5 1.473 Gd 4.8 8.8 4.5 8.6 6.0 5.675 Tb 0.9 1.2 0.8 1.4 1.0 0.95 Dy 4.9 6.8 4.6 7.7 5.4 5.2 Ho 1.0 1.3 0.9 1.5 1.1 1.05 Er 3.3 3.9 2.8 4.3 3.3 3.075 Tm 0.5 0.6 0.4 0.7 0.5 0.475 Yb 3.3 3.3 2.7 4.0 3.2 2.95 Lu 0.5 0.5 0.4 0.6 0.5 0.448 Note: N.D.—not determined; SN—averaged values from data from the Série Negra Ediacaran sediments of the Ossa-Morena zone (from Pereira et al., 2006).
strong shearing developed under very-low-grade metamorphism (e.g., Ribeiro, 1983; Carvalhosa and Zbyszewski, 1994; Silva et al., 2003; Chichorro et al., 2004). A kilometer-scale fan-like geometry centered on a syncline bounded by high-strain zones with moderately to steeply dipping foliation and weakly to moderately dipping stretching lineation characterizes the Montemor-o-Novo shear zone complex structural pattern (Figs. 3–5). Northeast steeply dipping mylonitic fabric is dominant in this tectonic unit, which includes a northern
syncline centered on basic rocks, bounded by high-strain shear zones, and unconformably preserving on top a weakly metamorphosed and gently deformed sedimentary and volcanic-sedimentary sequence (Cabrela Formation). Metamorphism with growth of medium-temperature/highpressure (550 °C, 11 kbar) mineral assemblages—garnet + omphacite + glaucophane (crossite) + quartz (Safira eclogites; e.g., Pedro, 1996; Leal et al., 1997; Leal, 2001)—is recognized within the high-strain shear zone that limits the main syncline in the southwest (see Figs. 1 and 3 for location). These high-pressure rocks occur along meter-scale boudins, surrounded by felsic gneisses and amphibolites that show a well-developed mylonitic fabric with synkinematic growth of feldspar + garnet + quartz + biotite and barroisite + white mica + plagioclase, respectively, recording a subsequent retrogression under medium-low pressure and medium-low temperature metamorphic conditions (amphibolitic–greenschist-facies). In this study, a metamorphic peak of medium temperature–medium pressure was identified on the Biscaia garnet-rich amphibolites (sample BSC-2), defined by decimeter-scale straight outcrops of foliated rocks surrounded by paragneisses from the Serie Negra at a distance of a few kilometers to the southeast of the Safira outcrop (Fig. 4). Garnet subhedral crystals (Fig. 6A–C and Table 3) show a core-rim zoning with an increase of mainly Mg and also Fe, decrease of Mn and Fe/(Fe + Mg) ratio, while Ca remains unchanged or with a slight decrease. P-T data obtained using calibrations of Garnet-Amphibole-Plagioclase (Kohn and Spear, 1990) and Garnet-Hornblende (Powell, 1985; Perchuck, 1991) revealed a pressure and temperature increase from the core (400–425 °C, 5–6.5 kbar) to the rim (450–600 °C, 6–7 kbar) during crystal growth (Fig. 6F). These porphyroblasts are surrounded by aligned crystals of hornblende + plagioclase + quartz + biotite, defining a well-developed planar fabric. Mineral chemistry of calcic hornblendes (Fe-hornblendes; Fig. 6E) used together with plagioclase (An24–37; Fig. 6D and Table 1A, B) suggests that these amphibolites have equilibrated during medium- to low-pressure and medium-temperature metamorphic conditions (Brown, 1974; Raase, 1974; Spear, 1981; Fig. 7C). Associated paragneisses of Biscaia (BSC-1) contain albite porphyroblasts with aligned and folded inclusions trails of unknown composition (Fig. 6B). These plagioclase porphyroblasts show asymmetrical strain shadows with crystallization continuing to the surrounded foliation, suggesting a synkinematic growth from albite to oligoclase (Fig. 6D) during progressive deformation in a noncoaxial regime of deformation. The surrounding foliation is defined by the alignment of feldspar + quartz + muscovite + biotite (amphibolitic facies) and also minor chlorite and sericite (greenschist facies), and by the development of shear-sense criteria revealing a sinistral movement along a stretching lineation moderately to weakly plunging to the northwest. Within the main syncline, mainly completed with amphibolites (Figs. 3 and 4), it is also possible to observe a prograde metamorphism developed under lower-pressure conditions. The structure is defined by the superposition of two planar fabrics as
Crustal growth and deformational processes in the northern Gondwana margin
339
TABLE 2. REPRESENTATIVE ELECTRON MICROPROBE ANALYSIS AND STRUCTURAL FORMULAS OF FELDSPAR FROM BSC-1 Sample BSC-1 BSC-1 BSC-1 BSC-1 BSC-1 BSC-1 BSC-1 BSC-1 BSC-1 BSC-1 BSC-1 Analysis P1 P2 P3 P4 P5 P6 P7 P8 P9 P10 P11 (wt%) SiO2 68.48 68.52 67.73 68.57 68.94 65.23 63.15 65.24 62.4 69.36 64.55 TiO2 0.04 0.01 0.0 0.0 0.0 0.0 0.0 0.0 0.01 0.02 0.0 Al2O3 20.33 19.48 19.58 20.32 19.73 22.54 23.26 23.01 23.77 19.7 22.94 Fe2O3 0.03 0.0 0.0 0.01 0.0 0.12 0.09 0.03 0.31 0.0 0.0 MgO 0.0 0.0 0.01 0.01 0.0 0.0 0.03 0.0 0.13 0.0 0.0 CaO 0.12 0.19 0.42 0.24 0.42 3.17 4.0 4.04 5.01 0.34 3.55 BaO 0.04 0.01 0.0 0.0 0.0 0.0 0.13 0.0 0.0 0.01 0.0 Na2O 10.69 10.5 10.36 10.54 10.16 8.89 8.56 8.37 7.45 10.33 8.44 K2O 0.11 0.12 0.12 0.13 0.1 0.09 0.08 0.1 0.2 0.15 0.13 Total 99.84 98.83 98.22 99.83 99.35 100.05 99.29 100.78 99.28 99.9 99.62 (ppm) Si 2984.0 3013.0 2999.0 2987.0 3012.0 2859.0 2803.0 2841.0 2772.0 3015.0 2842.0 Ti 0.001 0.0 0.0 0.0 0.0 0.0 0.0 0.0 0.0 0.0 0.0 IV Al/Al 1044.0 1009 1022.0 1043.0 1016.0 1164.0 1217.0 1181.0 1244.0 1009.0 1190.0 VI Al 0.0 0.0 0.0 0.0 0.0 0.0 0.0 0.0 0.0 0.0 0.0 3+ Fe 0.001 0.0 0.0 0.0 0.0 0.004 0.003 0.001 0.01 0.0 0.0 Mg 0.0 0.0 0.001 0.0 0.0 0.0 0.002 0.0 0.008 0.0 0.0 Ca 0.006 0.009 0.02 0.011 0.019 0.149 0.19 0.188 0.239 0.016 0.167 Ba 0.001 0.0 0.0 0.0 0.0 0.0 0.002 0.0 0.0 0.0 0.0 Na 0.903 0.895 0.889 0.89 0.861 0.756 0.736 0.707 0.642 0.871 0.721 K 0.006 0.007 0.007 0.007 0.005 0.005 0.004 0.005 0.011 0.008 0.007 Ab 98,648.0 98,282.0 97,069.0 97,953.0 97,179.0 83,081.0 78,912.0 78,494.0 71,953.0 97,307.0 80,497.0 An 0.619 0.975 2195.0 1254.0 2200.0 16,345.0 20,394.0 20,917.0 26,758.0 1751.0 18,709.0 Or 0.667 0.727 0.736 0.788 0.62 0.569 0.455 0.588 1289 0.931 0.793
Sample Analysis (wt%) SiO2 TiO2 Al2O3 Cr2O3 FeOt MnO MgO CaO Na2O K2O NiO Total (ppm) Si Al Ti Cr ++ Fe Mn Mg Ni Ca Na K
TABLE 3. REPRESENTATIVE ELECTRON MICROPROBE ANALYSIS AND STRUCTURAL FORMULAS OF AMPHIBOLE FROM BSC-2, SCD, REL, ARL-1, AND BDV-2 BSC-2 BSC-2 BSC-2 BSC-2 BSC-2 SCD SCD SCD SCD SCD SCD 286 293 298 301 304 323 core 324 rim 327 core 328 rim 330 core 331 rim 46.089 0.427 9.385 0.0 20.611 0.402 7.992 11.414 1.097 0.373 0.07 97.86
44.358 0.469 12.041 0.0 20.604 0.497 8.07 10.024 1.32 0.39 0.0 97.773
49.544 0.359 5.578 0.076 19.571 0.409 10.318 11.568 0.7 0.178 0.042 98.343
43.507 0.485 13.307 0.034 20.239 0.445 7.256 10.723 1.495 0.428 0.0 97.919
46.947 0.427 8.502 0.007 19.786 0.223 9.151 11.671 0.977 0.239 0.0 97.93
53.782 0.128 2.116 0.077 11.223 0.297 16.629 12.558 0.303 0.105 0.084 97.302
44.606 0.794 9.292 0.0 17.485 0.426 11.107 11.775 1.289 1.032 0.051 97.858
46.386 0.679 8.361 0.0 16.432 0.292 12.034 12.057 1.097 0.791 0.111 98.24
44.964 0.697 9.468 0.096 16.669 0.336 11.202 12.278 1.095 1.044 0.0 97.848
41.327 0.869 12.056 0.02 18.51 0.529 9.318 11.983 1.355 1.624 0.0 97.591
42.123 0.575 11.471 0.132 18.039 0.318 9.89 12.041 1.39 1.191 0.042 97.212
6.951 1.669 0.048 0.000 2.600 0.051 1.796 0.008 1.845 0.321 0.072
6.683 2.139 0.053 0.000 2.596 0.063 1.812 0.000 1.618 0.386 0.075
7.357 0.976 0.040 0.009 2.430 0.051 2.283 0.005 1.841 0.202 0.034
6.555 2.364 0.055 0.004 2.550 0.057 1.629 0.000 1.731 0.437 0.082
7.032 1.501 0.048 0.001 2.478 0.028 2.043 0.000 1.873 0.284 0.046
7.736 0.359 0.014 0.009 1.350 0.036 3.565 0.010 1.936 0.085 0.019
6.708 1.647 0.090 0.000 2.199 0.054 2.489 0.006 1.897 0.376 0.198
6.880 1.462 0.076 0.000 2.038 0.037 2.660 0.013 1.916 0.315 0.150
6.731 1.671 0.078 0.011 2.087 0.043 2.499 0.000 1.969 0.318 0.199
6.324 2.175 0.100 0.002 2.369 0.069 2.125 0.000 1.965 0.402 0.317
6.429 2.064 0.066 0.016 2.303 0.041 2.250 0.005 1.969 0.411 0.232 Continued
340
Sample Analysis (wt%) SiO2 TiO2 Al2O3 Cr2O3 FeOt MnO MgO CaO Na2O K2O NiO Total (ppm) Si Al Ti Cr ++ Fe Mn Mg Ni Ca Na K
Pereira et al. TABLE 3. REPRESENTATIVE ELECTRON MICROPROBE ANALYSIS AND STRUCTURAL FORMULAS OF AMPHIBOLE FROM BSC-2, SCD, REL, ARL-1, AND BDV-2 (continued) BDV-2 BDV-2 BDV-2 BDV-2 BDV-2 BDV-2 ARL-1 ARL-1 ARL-1 ARL-1 ARL-1 1 2 4 5 6 8 1 core 2 rim 3 core 5 core 6 core 52.12 0.39 5.44 0.01 12.52 0.27 16.5 10.43 0.87 0.06 0.03 100.8 7.408 0.911 0.042 0.001 1.488 0.033 3.495 0.003 1.588 0.240 0.011
44.76 0.98 12.23 0.07 14.65 0.41 12.11 10.47 2.44 0.07 0.1 100.37 6.555 2.111 0.108 0.008 1.794 0.051 2.643 0.012 1.643 0.693 0.013
44.41 0.97 12.92 0.0 13.75 0.37 11.87 11.06 2.41 0.11 0.02 99.94 6.510 2.233 0.107 0.000 1.686 0.046 2.593 0.002 1.737 0.685 0.021
45.52 0.79 12.58 0.08 13.33 0.42 12.29 11.18 2.24 0.12 0.0 100.64 6.601 2.151 0.086 0.009 1.617 0.052 2.656 0.000 1.737 0.630 0.022
45.67 0.9 12.05 0.02 13.83 0.27 12.53 10.78 2.33 0.06 0.05 100.6
45.56 1.01 12.15 0.0 12.95 0.25 12.52 11.41 2.23 0.1 0.14 100.38
6.632 2.063 0.098 0.002 1.680 0.033 2.712 0.006 1.677 0.656 0.011
47.1 0.69 8.78 0.04 15.62 0.34 12.12 11.68 1.19 0.32 0.12 100.05
6.616 2.080 0.110 0.000 1.573 0.031 2.709 0.016 1.775 0.628 0.019
47.36 0.7 8.7 0.0 14.92 0.34 12.18 11.96 1.15 0.34 0.09 99.79
6.936 1.524 0.076 0.005 1.924 0.042 2.660 0.014 1.843 0.340 0.060
6.969 1.509 0.077 0.000 1.836 0.042 2.671 0.011 1.886 0.328 0.064
47.86 0.43 9.31 0.04 15.71 0.21 11.8 10.88 0.95 0.3 0.05 99.61 7.030 1.612 0.047 0.005 1.930 0.026 2.583 0.006 1.712 0.271 0.056
47.81 0.58 8.69 0.01 15.01 0.29 12.64 11.84 1.12 0.31 0.0 100.37 6.983 1.496 0.064 0.001 1.833 0.036 2.751 0.000 1.853 0.317 0.058
48.77 0.6 7.79 0.0 15.04 0.42 13.04 11.89 0.92 0.25 0.1 100.92 7.079 1.333 0.065 0.000 1.826 0.052 2.821 0.012 1.849 0.259 0.046 Continued
TABLE 3. REPRESENTATIVE ELECTRON MICROPROBE ANALYSIS AND STRUCTURAL FORMULAS OF AMPHIBOLE FROM BSC-2, SCD, REL, ARL-1, AND BDV-2 (continued) ARL-1 ARL-1 ARL-1 REL REL REL REL REL REL REL 7 rim 8 core 11 core 1 core 2 rim 3 core 4 rim 5 core 6 rim 7 core
Sample Analysis (wt%) SiO2 47.22 47.18 TiO2 0.63 0.53 Al2O3 8.87 8.83 Cr2O3 0.0 0.07 FeOt 15.32 15.24 MnO 0.32 0.25 MgO 12.02 12.35 CaO 11.68 11.75 Na2O 1.12 1.12 K2O 0.31 0.31 NiO 0.0 0.0 Total 99.54 99.69 (ppm) Si 6.968 6.952 Al 1.543 1.534 Ti 0.070 0.059 Cr 0.000 0.008 ++ Fe 1.891 1.878 Mn 0.040 0.031 Mg 2.643 2.712 Ni 0.000 0.000 Ca 1.847 1.855 Na 0.320 0.320 K 0.058 0.058 Note: FeOt—total iron as FeO.
48.25 0.71 8.22 0.01 15.58 0.24 12.48 11.56 1.12 0.23 0.06 100.54 7.041 1.414 0.078 0.001 1.901 0.030 2.714 0.007 1.808 0.317 0.043
56.75 0. 0.68 0.02 10.38 0.31 17.43 12.94 0.06 0.0 0.0 100.71 7.980 0.113 0.000 0.002 1.221 0.037 3.653 0.000 1.950 0.016 0.000
45.32 0.42 10.37 0.02 16.36 0.22 11.05 12.24 1.37 0.28 0.0 99.67 6.743 1.819 0.047 0.002 2.036 0.028 2.450 0.000 1.951 0.395 0.053
55.73 0.02 1.25 0.03 10.7 0.28 17.01 12.79 0.13 0.02 0.07 100.15 7.904 0.209 0.002 0.003 1.269 0.034 3.595 0.008 1.944 0.036 0.004
45.08 0.51 10.29 0.03 16.45 0.29 10.62 12.39 1.36 0.24 0.0 99.28 6.744 1.815 0.057 0.004 2.058 0.037 2.368 0.000 1.986 0.395 0.046
55.5 0.06 1.34 0.02 10.87 0.23 17.16 12.87 0.11 0.04 0.0 100.33 7.867 0.224 0.006 0.002 1.289 0.028 3.625 0.000 1.955 0.030 0.007
45.33 0.47 10.37 0.03 16.85 0.29 10.91 12.22 1.4 0.23 0.1 100.24 6.726 1.814 0.052 0.004 2.091 0.036 2.412 0.012 1.943 0.403 0.044
45.1 0.4 10.19 0. 17.81 0.23 9.98 12.24 1.4 0.21 0.08 99.65 6.760 1.801 0.045 0.000 2.232 0.029 2.229 0.010 1.966 0.407 0.040
Crustal growth and deformational processes in the northern Gondwana margin
Sample Analysis (wt%) CaO TiO2 Cr2O3 MnO FeOt SiO2 Al2O3 MgO Total (ppm) Si Ti IV Al VI Al Cr 3+ Fe 2+ Fe Mg Mn Ca Alm Pyr Gros Sp XFe
341
TABLE 4. REPRESENTATIVE ELECTRON MICROPROBE ANALYSIS AND STRUCTURAL FORMULAS OF GARNET FROM BSC-2 BSC-2 BSC-2 BSC-2 BSC-2 BSC-2 BSC-2 BSC-2 BSC-2 BSC-2 Grt1-P1 Grt1-P2 Grt1-P3 Grt1-P4 Grt1-P5 Grt2-P1 Grt2-P2 Grt2-P3 Grt2-P4 10.306 0.102 0.0 3.53 27.019 37.568 21.107 0.531 100.253
10.337 0.18 0.013 3.033 27.107 37.79 21.335 0.497 100.348
10.649 0.153 0.06 1.087 28.148 37.615 21.194 0.625 99.561
10.043 0.157 0.039 0.746 30.465 37.542 20.568 0.809 100.429
9.98 0.052 0.066 0.595 31.18 37.81 21.016 1.011 101.728
9.898 0.123 0.007 3.285 26.881 37.54 21.32 0.592 99.804
10.413 0.198 0.034 4.415 25.439 37.895 20.922 0.416 99.766
10.154 0.152 0.073 2.741 27.753 37.47 21.205 0.584 100.156
3.003 0.006 0.0 1.989 0.0 0.004 1.802 0.063 0.239 0.882 0.603 0.021 0.291 0.080 0.966
3.016 0.010 0.0 2.007 0.001 0.0 1.809 0.059 0.205 0.884 0.611 0.019 0.294 0.069 0.968
3.020 0.009 0.0 2.006 0.003 0.0 1.890 0.074 0.073 0.916 0.639 0.025 0.305 0.025 0.961
2.999 0.009 0.0 1.937 0.002 0.050 1.984 0.096 0.050 0.859 0.663 0.032 0.264 0.016 0.953
2.980 0.003 0.019 1.933 0.004 0.059 1.996 0.118 0.039 0.842 0.665 0.039 0.256 0.013 0.943
3.013 0.007 0.0 2.017 0.0 0.0 1.804 0.070 0.223 0.851 0.611 0.024 0.285 0.075 0.962
3.044 0.011 0.0 1.981 0.002 0.004 1.704 0.049 0.300 0.896 0.577 0.016 0.297 0.101 0.971
2.998 0.009 0.001 1.999 0.004 0.0 1.857 0.069 0.185 0.870 0.622 0.023 0.286 0.062 0.963
9.719 0.155 0.0 1.534 28.7 37.838 20.969 0.648 99.603 3.044 0.009 0.0 1.989 0.0 0.001 1.929 0.077 0.104 0.837 0.654 0.026 0.280 0.0354 0.9612
BSC-2 Grt2-P5 9.972 0.187 0.0 0.733 29.287 37.981 21.054 0.814 100.041 3.039 0.011 0.0 1.986 0.0 0.002 1.957 0.097 0.049 0.854 0.661 0.032 0.283 0.016 0.952
Note: Alm—almandine; Gros—grossular; Pyr—pyrope; Sp—spessartite; XFe—XFe (iron).
a result of a sequence of deformation stages related to progressive deformation. The texture of the amphibolites is defined by the alignment of amphiboles and plagioclase (Fig. 7A, B), while associated metasediments present a planar fabric with growth of andalusite + biotite + quartz + feldspar ± garnet. Colorless cores characterize aligned green amphiboles that define the foliation. These calcic amphiboles show an increase of Mg content (Fig. 7A, B and Table 2) from cores (actinolite) to rims (Mghornblende) caused by an increase in temperature. The plagioclase also suggests a synkinematic growth from albite to andesine due to prograde metamorphism. The composition of these calcic amphiboles has AlIV-Si relationships and small NaM4/AlIV ratios (Fig. 7C), indicative of equilibration during low-pressure metamorphic conditions. At the northern limit of the main syncline another highstrain shear zone was defined just along the southern border of the Évora high-grade metamorphic terrains (Silveiras–Boa Fé; Figs. 1 and 4). This shear zone shows a characteristic isograd pattern, with tight spacing passing from the biotite to the sillimanite zone within a distance of a few hundred meters, suggesting a strong thermal gradient. Along a well-developed northwest–
southeast-trending and steeply dipping to southwest mylonitic foliation (Fig. 4), a gently northeast-dipping stretching lineation indicates sinistral shear-sense criteria of movement, based on CS structures, asymmetric pressure-shadow tails, and extensional shear cleavage (Chichorro et al., 2003, 2004). The incidence of the biotite + andalusite + fibrolitic sillimanite + quartz + plagioclase + cordierite mineral assemblage is related to the first stages of leucosome differentiation and seems to be clearly associated with the highest strain gradients characterized by well-developed mylonitic textures. The presence of cordierite + sillimanite associated with feldspar blastesis (Chichorro, 2006), as well as the lack of garnet in pelitic rocks, suggests a prograde evolution under low-pressure conditions (3.5–4.5 kbar), reaching temperatures high enough to cause dehydration melting reactions (650–750 °C). The biotite + andalusite + sillimanite + cordierite mineral assemblage starts with the abrupt appearance of brownred biotite followed by xenoblastic andalusite and the fibrolitic sillimanite. Andalusite is residual (xenoblastic), and cordierite increases in banded paragneisses. The intergrowth between biotite and fibrolite-sillimanite is common. The sillimanite occurs as strong elongated clusters of fibrolite needles parallel to the
V
V
V V
V V
V
V V
V
V
V
V
V
V
V
Foliated granodiorites
Igneous (basic-dominated)-sedimentary complex (includes the Carvalhal Formation Upper Cambrian-Lower Ordovician)
Unconformity
Cabrela Formation Lower Carboniferous (* Pedreira de Engenharia Olistostrome Middle Devonian)
Serie Negra (Escoural Formation) Ediacaran
Unconformity
Igneous (felsic-dominated)-sedimentary V complex (includes the Monfurado Formation Lower-Middle Cambrian)
V V V V V V V V V
V
V
V
+
V
+ +
+
V
V
+
+
V
X
Foliated tonalites, granodiorites, gabbro-diorites ca. 320 Ma
Granites and Porphyritic granodiorites ca. 310 Ma
Gneiss-migmatitic Complex
VX
X
+
+ +
X
+
+
Évora High-grade Metamorphic Terrains
V
V V
V
V
X
V
V
+ +
X
X
Foliated tonalites and granodiorites
Serie Negra (Águas de Peixe Formation) Ediacaran
Unconformity
Igneous (felsic-dominated)-sedimentary complex (include the Dolomitic Formation Lower-Middle Cambrian)
Igneous (basic-dominated)-sedimentary complex (includes the Xistos de Moura Formation Upper Cambrian-Lower Ordovician?)
X
Granites and Porphyritic granodiorites
Évora Medium-grade Metamorphic Terrains
Figure 2. Interpretative tectonostratigraphic columns for the Neoproterozoic-Paleozoic rocks. Modifed from Carvalhosa (1983, 1999), Oliveira et al. (1991), Carvalhosa and Zbyszewski (1994), and Pereira et al. (2003a).
*
V V V
Montemor-o-Novo Shear Zone
342 Pereira et al.
Crustal growth and deformational processes in the northern Gondwana margin Safira
343
Cabrela
MNSZ
N40ºW
A
B N35ºE EMMT
Silveiras MNSZ
C
D
Biscaia MNSZ
EHMT
Montemor
N30ºE
E
F S. Conde
Monfurado
EHMT N60ºE
MNSZ
G
H Valverde
Évora
EMMT
EHMT
EHMT
EMMT N70ºE
+
I
+
+ + +
+
+
+ +
+ +
+
+ +
+
Transcurrent faults Normal faults
+
+
+
+ +
+
+
+ +
+
+
+
+
Tonalites (Carboniferous) + + Serie Negra (Ediacaran)
+
+
+
Granites (Carboniferous)
Igneous(felsic)-sedimentary complex (Lower-Middle Cambrian)
J
Cabrela Formation (Lower Carboniferous) Igneous(basic)-sedimentary complex (Upper Cambrian-Lower Ordovician)
Figure 3. Interpretative cross-sections (marked A–J) showing the Évora Massif main structural characteristics. See Figure 1 for location of crosssections. EHMT—Évora high-grade metamorphic terrains; EMMT—Évora medium-grade metamorphic terrains; MNSZ—Montemor-o-Novo shear zone.
stretching lineation or, more rarely, as isolated fibrolitic clots intergrown with quartz and K-feldspar. Paragneisses exhibit thin and periodic venules, subparallel to the main foliation and essentially shaped by quartz and small amounts of plagioclase. Their borders are regular and make contact with the host rock by cordierite-rich bands. Eye-shaped nodules of andalusite + biotite + cordierite with the xenoblastic andalusite in equilibrium with red biotite + quartz, and rounded by cordierite extensively altered to pinite, are characteristic. These nodules are rounded by intergrowths of red biotite, fibrous sillimanite, and quartz. Évora High-Grade Metamorphic Terrains The Évora high-grade metamorphic terrains, with an average width of 15–20 km, are formed by a structurally complex assemblage of variably sheared migmatites and gneisses associated to variably strained N300° to 310°–trending granodiorites, tonalites, gabbros, diorites (Moita et al., 2005a), and andesitic
dikes, all intruded by an undeformed suite of granites and porphyritic granodiorites (Figs. 1–3). The high-grade metamorphic rocks correspond mainly to biotite-rich paragneisses, cordieritebearing migmatites, and banded leucogneisses and may include restites of black metacherts, metapelites, metagraywackes, calcsilicate marbles, and amphibolites (from the Neoproterozoic and Lower Paleozoic sequences; Pereira and Silva, 2002). This unit is bounded to the south by the Montemor-o-Novo shear zone and toward the north by the Évora medium-grade metamorphic terrains, both essentially composed of strongly sheared mica schists, quartz-phyllites, amphibolites, and felsic gneisses derived from Neoproterozoic and Lower Paleozoic protoliths. The highest temperatures were reached within the Évora high-grade metamorphic terrains, which represent a dome-like geometry made of anatexis granitoids and migmatitic orthoand paragneisses, which has a mineral assemblage of high amphibolitic facies and is transitional between the amphibolite and granulite facies. The mineral assemblages developed in these
N
N
Safira quartz-feldspar gneiss
N
Qtz ribbons
Biscaia paragneiss
N
5
Qtz ribbons
32
+
+ +
+
8
+
+
+
24
1
+
50
+
+
4
+
+
Pl porphyroclast Qtz tails and ribbons
Foliation with Ms + Bt + Qtz
+
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+
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6
+
2 mm
Top to NW
+
5
60 4
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+
7 80 80
+ +
Montemor-o-Novo 70
+
80
+
74 74
+
Escoural quartz-feldspar gneiss
Sinistral C shear planes
FK blasts with graphitic folded fabric
2 mm
Foliation with Ky + Bt + Qtz
+
50 18
+
+
+
EMMT
8º 15’
6
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+
+
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+
+ +
+
+
+
+
+ +
+
+
+
+
+
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+
+
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86+
9 86 80
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8015
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70 24
80
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Valverde
+
+
+
20 38 +
Arraiolos
74
+
+
+
+
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+
Évora
+
+
+
+
56 + 56
+
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V
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+
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10
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+
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+
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50
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20
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+
80
+
+
+ + +
80 40
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3 Km
+
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+
+
54
4
24
+
+
+
Qtz ribbons
1 mm
Top to NW
N
Pl
6 22 10
10 10
20
38 38
38º 40’
Crd
7
Foliation with Pl + Hb + Qtz
Sinistral C’extensional shear planes
N
Top to NW
N N
Foliation with Bt + Sil
Top to SE
And + Qtz + Bt
Sinistral C’ extensional shear bands
Carvalhal amphibolite
0.5 mm
Dextral C’ extensional shear bands
Top to E
N
Casas Novas cordierite-rich paragneiss
Aggregates of HB with Cpx
0.5 mm
8
1 mm
Foliation with Hb + Pl + Qtz
Valverde amphibolite
1 mm
9
Oblique foliation on quartz ribbons
And
Top to N
N
Sil + Bt + Qtz
Évora and aluzite-rich paragneiss
Hb blasts with cores of Act
Escoural biotite-rich quartz-feldspar gneiss
+
+
+
+
+
Conjugate shear planes Fd porphyroclast
4
+
+
+
FK blasts with graphitic folded fabric
+
8
16
+
+
N
Sinistral C’ extensional shear bands
10
Figure 4. Simplified structural sketch map of the Évora Massif. Detailed schemes of thin-sections with microstructures and metamorphic assemblages from (1) to (10) to depict the S-L fabric variable distribution and orientation. Equal area lower hemisphere stereographic projections show the stretching lineation (black circle)–foliation (black line) relationship. Act—actinolite; Bt—biotite; Ch—chlorite; Cpx—clinopyroxene; Crd—corderite; EHMT—Évora high-grade metamorphic terrains; EMMT—Évora medium-grade metamorphic terrains; Ep—epidote; Fd—feldspar; Fk—K-feldspar; Grt—garnet; HB—hornblende; Ky—kyanite; Ms—muscovite; MNSZ—Montemor-o-Novo shear zone; Op—opaque minerals; Pl—plagioclase; PTFZ—Porto-Tomar fault zone; Qz—quartz.
Foliation with Fd porphyroclast Ms + Qtz
Sinistral C’extensional shear planes
Top to NW
4
PTFZ
30 30
+
EHMT
Cabrela
0.5 mm 0.5 mm
Grt + Op + Gc
Dextral C’extensional shear planes
Fd blasts with tails of Bt + Ch
MNSZ
Dextral C’extensional shear planes
Fd blasts Top to NE with Bt inclusions Dextral C’extensional shear planes
3
0.5 mm
Foliation with Ph + Bt + Qtz
Pl blasts Foliation with Pl + Hb + Qtz
Safira biotite-rich quartz-feldspar gneiss
Top to NE
2
0.5 mm
Dextral C shear planes
Ep+Ch tails
Sinistral C’extensional shear planes
Hb blasts with cores of Act
Carvalhal amphibolite
Top to SE
1
344 Pereira et al.
Crustal growth and deformational processes in the northern Gondwana margin
345
Évora Medium-grade Metamorphic Terrains Local dip-slip stretching lineation N
Évora High-grade Metamorphic Terrains
20
Dominant moderately dipping foliation
38
56
24
Montemor-o-Novo Shear Zone
20 30
48 Local dip-slip stretching lineation
8 15 16 10
40
Local flat-lying foliation
4 24
20
10
6
4
4
Dominant flat-lying foliation
86 10
Local dip-slip stretching lineation Dominant steeply dipping foliation
6
24
Transcurrent movement Major detachments Normal fault
Figure 5. Idealized block diagram illustrating fabrics and kinematics in the Évora Massif. Gray shades represent the gneiss-migmatitic complex and associated granitoids, the black arrows the stretching lineation, and the white arrows the sense of movement. Note the spatial variation of the S-L fabric: in some sectors flat-lying foliation dominates (Évora medium-grade metamorphic terrains), whereas in others, steeply dipping foliation exists (Montemor-o-Novo shear zone).
rocks constitute quartz + biotite + plagioclase + K-feldspar + sillimanite + cordierite ± andalusite ± garnet (scarce) (e.g., Carvalhosa and Zbyszewski, 1994) and were associated with progressive partial melting with dehydration reactions of muscovite and biotite. The subsequent decrease of temperature produced lowgrade assemblages and destabilization of cordierite, biotite, and feldspar, with the growth of chlorite and albite. The intrusion of tonalites took place while the Variscan migmatization was in progress, causing local partial melting and mixing of melts. There is so far no field evidence to support the existence of an older (Cadomian) high-temperature metamorphic event. The northwest–southeast-trending steep foliation, variable thick veins of leucosome, shear bands, and asymmetric isoclinal folding (often rootless) are the most prominent features of these high-grade rocks, which include several meter- to decimeterscale boudins of amphibolites (Lower Paleozoic), paragneisses, and black metacherts (Neoproterozoic). Within this tectonic unit the gently dipping stretching lineation marked by the alignment
of biotite or sillimanite progressively loses intensity and becomes more widely spaced as a consequence of well-developed granoblastic textures and the heterogeneous distribution of deformation. Rotation of previous structures caused by progressive sinistral shearing allowed the dispersion of fold-axis orientations, the appearance of an oblique foliation (Pereira and Silva, 2002), and the creation of dilatant surfaces for the leucosomes flow (surfaces subparallel to foliation, boudins tension tails, boudins intrafractures, C′-type shear bands, and fold hinges). Évora Medium-Grade Metamorphic Terrains The Évora medium-grade metamorphic terrains represent a 10- to 20-km-wide band of rocks deformed under amphibolitic-facies conditions that extends for ~35 km from Arraiolos to Valverde as the northernmost limit of the Évora high-grade metamorphic terrains. The stratigraphic sequence (Figs. 2 and 3) is mainly composed by Middle-Upper Cambrian–Ordovician(?)
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Amphibole Plagioclase
BSC-2
P5 Rim P4
P5 Rim P3 P4 P2 P1 Core
A
%Or
P3
BSC-1
Microcline
P2 P1 Core
Blast tails (P5, P6, P11) Blast tails (P7, P8, P9)
0.2 mm Garnet 1
Blast with an internal folded fabric (P1,P2,P3,P4,P10)
Perthitic Microcline
K-Feldspar Garnet 2
P7 P9
P8
P6
P5 P4
P3
BSC-2 Garnet corona
BSC-1
P2 P1
Matrix
Myrmekite Albite
P10
P11
%Ab
%An Oligoclase Andesine Labradorite Bytownite Anorthite
D
0.2 mm
B Oligoclase
Micas Feldspar + Quartz
Albite
0.980
X Fe
Calcic Amphiboles [(Na+K)A < 0.5; Ti <0.5] 1.0 0.9 Actinolitic Hornblende 0.8 Mg- Hornblende 0.7 Actinolite 0.6 0.5 0.4 Fe- Hornblende 0.3 0.2 Fe-Actinolite Fe-Actinolitic Hornblende 0.1 0.0 8.00 7.50 7.00 6.50 Tremolitic Hornblende
Tremolite
Mg/(Mg+Fe2+)
BSC-2
0.970 0.960
Garnet 2
0.950
Garnet 1
0.940 0.930
Alm
Prp
3.0 2.0
14
1.0
Safira eclogites
12
0.0 40.0
Pressure kbar
Grs
30.0 20.0 10.0 0.0 30.0
10 8
Safira gneisses BSC-2 Grt Rim
BSC-2 Grt Core
6 4
Casas Novas gneisses and migmatites
Ky
Sps
Sil
2
20.0
E
Water-saturated melting for granitic compositions
0.920 80.0 70.0 60.0 50.0 40.0 30.0 20.0 10.0 0.0 4.0
And SCD amphibolites
C
10.0 0.0
P1 Core
P2
P3
P4
400
600 500 700 Temperature ºC
800
F
P5 Rim
Figure 6. Thin-section schemes of the Biscaia (A) garnet-rich amphibolites (sample BSC-2; detail of a fractured garnet surrounded by a K-feldspar rim, within an amphibole + plagioclase matrix) and (B) paragneisses (BSC-1; albite porphyroblast with an internal folded fabric of unknown composition and with oligoclase tails); P1–P11 are the locations for the mineral chemistry plots. (C) Rim-to-core compositional profiles in garnet from the BSC-2. Alm—almandine; Grs—grossular; Prp—pyrope; Sps—spessartite; X Fe—iron. (D) Or-Ab-An plot for average compositions of feldspars from BSC-1 and BSC-2. Ab—albite; An—anorthite; Or—orthoclase. (E) Compositions of iron-magnesian amphiboles of BSC-2, with fields from Leake (1978). And—andalusite; Grt—garnet; Ky—kyanite; SCD—sample name; Sil—sillimanite. (F) Pressure-temperature path followed by the Biscaia garnet-rich amphibolites and Serra do Conde amphibolites (this study), the Casas Novas gneisses and migmatites (data from Chichorro et al., 2004), and the Safira gneisses and eclogites (data from Pedro, 1996; Leal, 2001).
%Or
A
Microcline
Crustal growth and deformational processes in the northern Gondwana margin
347
SCD
Perthitic Microcline
Rims
Rim
Albite
%Ab
%An Oligoclase Andesine Labradorite Bytownite Anorthite
Cores
Calcic Amphiboles [(Na+K)A < 0.5; Ti <0.5]
Core
1.0 0.9 0.8 Mg- Hornblende 0.7 Actinolite 0.6 Rims 0.5 0.4 Fe- Hornblende 0.3 0.2 Fe-Actinolite Fe-Actinolitic Hornblende 0.1 0.0 8.00 7.50 7.00 6.50 Tremolitic Hornblende
Mg/(Mg+Fe2+)
Tremolite
0.2 mm mm
B
Cores
Core
Rim
Actinolitic Hornblende
%Or Microcline
REL Perthitic Microcline
Albite
%Ab
%An Oligoclase Andesine Labradorite Bytownite Anorthite
Calcic Amphiboles [(Na+K)A < 0.5; Ti <0.5] Cores
1.0 0.9 0.8 Mg- Hornblende 0.7 Actinolite 0.6 Rims 0.5 0.4 Fe- Hornblende 0.3 0.2 Fe-Actinolite Fe-Actinolitic Hornblende 0.1 0.0 8.00 7.50 7.00 6.50 Tremolite
Actinolitic Hornblende
Mg/(Mg+Fe2+)
mm 0.2 mm
Tremolitic Hornblende
2.0
C
1.8 1.6
1.00
7 kbar
1.2
C 0º 49 ºC 35
1.4
72 5
Na M4
x. A lV HP HP
I (L eak e,
19
LP LP
78
0.4
6 kbar
1.0
5k
bar
0.2
5kbar
0.8 0.0
5 kbar
0.6
0.20
5.0
5.5
6.0
4 kbar
0.4
0.1
0.2
0.3
0.4
X Na (A) in amphibole
0.5
0.0
7.0
7.5
8.0
SCD
2 kbar
0.0
0.0
6.5 Si
3 kbar
0.2
0.00
)
0.6
1.2
0.40
1.0 0.8
ºC
0º
C
1.6
60
X Ab in plagioclase
5
0.60
Al VI
1.8
0.80
Ma
1.4
2.0
0.5
1.0 Al IV
1.5
BSC-2 REL BDV-2 ARL-1
Figure 7. (A, B) Photographs of thin-section of amphibolites from the Montemor-o-Novo shear zone showing calcic amphiboles with rim-to-core compositional variations. SCD and REL are sample names. Or-Ab-An plot for average compositions of feldspars from BSC-1 and BSC-2 and compositions of iron-magnesian amphiboles with fields from Leake (1978). (C) Semiquantitative temperature and pressure estimates for amphibole-plagioclase equilibria in amphibolites from the Évora medium-grade metamorphic terrains and the Montemor-o-Novo shear zone (after Brown, 1974; Raase, 1974; Spear, 1981). Ab—albite; An—anorthite; HP—high pressure; LP—low pressure; Or—orthoclase.
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igneous (basic-dominated)-sedimentary complex with mica schists, amphibolites, quartzites, and minor calc-silicate rocks (Xistos de Moura Formation). The sequence locally includes the Cadomian basement (Serie Negra; Águas de Peixe Formation) overlain by a Cambrian igneous (felsic-dominated)-sedimentary complex (Dolomitic Formation) with marbles, felsic gneisses and minor amphibolites, and/or calc-silicate rocks. Gabbros-diorites of Carboniferous age that show variable deformation intruded this Neoproterozoic–Lower Paleozoic stratigraphic sequence. The Evora medium-grade metamorphic terrains represent a tectonic unit (hangingwall) that is separated from the Évora high-grade metamorphic terrains (footwall) by a folded major detachment. Here the geometry is characterized by predominance of flat-lying planar fabrics (Figs. 4 and 5) and heterogeneous distribution of variable and weakly to moderately plunging fold axes (to the northwest or southeast) parallel to stretching lineation defined by amphiboles, biotites, or sillimanite (Fig. 4). Shear criteria are consistent with an orogen-parallel to slightly oblique sense of movement, but dip-slip lineation may locally exist (Fig. 5). The temperature of metamorphism associated with shearing was lower than that of the Évora high-grade metamorphic terrains, and the distribution of metamorphic isograds and deformation, along the limit between these two tectonic units, suggests an increase of temperature (with local partial melting) and shearing (local intensification of mylonitization) toward the southwest. Low-pressure/medium- to high-temperature metamorphism is characterized by the presence of several mineral assemblages on metasediments (e.g., Carvalhosa, 1999): quartz + K-feldspar ± plagioclase + biotite + sillimanite + garnet ± cordierite (sillimanite zone: mica schists close to the tonalitic intrusions), quartz + biotite + andalusite + plagioclase + muscovite (andalusite zone: mica schists), and quartz + plagioclase (oligoclase) + biotite + muscovite + andalusite + sillimanite ± garnet (scarce) (metapsammites). Andalusite porphyroblasts have tails with sillimanite growth and sometimes include an internal fabric made of opaque minerals and quartz. Sillimanite also occurs, intergrowing with quartz and feldspar along the foliation planes. Amphibolites occur, associated with metasediments. Strong development of mylonitization caused amphibole and plagioclase dynamic recrystallization, and the texture became fine-grained. Plagioclase, quartz, and amphibole are the most often observed minerals, but biotite, chlorite, epidote, pyroxene (locally found), sphene, and opaques also occur. These amphiboles are mostly Mg-hornblende (ARL-1) or Mg-hornblende and actinolitic hornblende (BDV-2; see Table 2) with low NaM4 values and AlIV-Si relationships, suggesting that they equilibrated at low pressure, whereas the semiquantitative temperature estimates from amphibole-plagioclase equilibrium of Spear (1981) indicate low- to medium-temperature conditions (Fig. 7C). The original gently dipping attitude of this planar structure was disturbed by the effect of doming, emplacement of late plutons, and later, movements along brittle faults. Rotated flat-lying foliation is observed at Valverde (Figs. 3 and 5), where a steeply dipping to vertical stretching lineation indicates, by several
shear-sense criteria (extensional cleavage, C-S structures, asymmetric pressure shadows, and sigmoidal systems), the exhumation of the western block (the migmatites from the Évora highgrade metamorphic terrains). MAIN EPISODES OF MAGMATISM Three main episodes of magmatism are recorded in the Évora Massif. The first one is preserved on Ediacaran sedimentary sequences made of immature graywackes that have resulted from a dismantled and reworked Cadomian continental magmatic arc (Pereira and Chichorro, 2004; Pereira et al., 2006). The second has Lower Paleozoic (Cambrian to Ordovician?) acid and basic volcanism and cogenetic plutonic rocks; precise crystallization ages for the uppermost complex are unknown and it is coeval with sedimentary sequences dominated by detrital and carbonate sediments (Pereira et al., 2005). Finally, the third episode of Carboniferous calc-alkaline volcanism and voluminous plutonism is composed mainly of tonalites, gabbros, diorites, and late orogenic granodiorites and granites (Moita et al., 2005a,b). In this section we present whole-rock geochemistry for major, trace, and rare earth element (REE) contents of thirty-six samples (see Fig. 8 for locations) analyzed by using the lithium metaborate/tetraborate fusion inductively coupled plasma mass spectrometry (ICP-MS) from ACTLABS, Canada. Table 5 shows the element composition of Lower (twenty-six samples) and Upper (ten samples) Paleozoic magmatic rocks. Immobile elements that produce cations with high charge/radius ratios (e.g., Pearce and Cann, 1973; Winchester and Floyd, 1977; Rollinson, 1993) were used to characterize the geochemistry of respective protoliths and thus to constrain hypotheses for igneous petrogenesis and the related geodynamic settings of Paleozoic magmatic rocks. This procedure was applied with the purpose of avoiding changes in the concentrations of the elements with low ionic potential caused by regional amphibolitic-facies metamorphism that affected Lower Paleozoic rocks. Eroded and Reworked Cadomian Continental Magmatic Arc The Ediacaran detrital rocks from the Ossa-Morena zone show a geochemical signature that essentially reflects the nature of an inherited continental margin source caused by the input of Cadomian detritus (e.g., Pereira and Chichorro, 2004; Pereira et al., 2006). In the Evora Massif, the Ediacaran sediments fall into the graywacke field according to their Al2O3, SiO2, Na2O, K2O, MgO, and Fe2O3 contents (Pettijohn et al., 1972; Blatt et al., 1980) and reveal low maturity, as reflected in their abundances of phyllosilicates and feldspar. The chondrite-normalized REE distribution patterns show a characteristic negative Eu anomaly and a clear enrichment of the light REE (LREE) relative to heavy REE (HREE; averaging high values of LaN/YbN = 14.2) similar to recent synorogenic volcanics from continental arcs (Fig. 9). The chondrite-normalized incompatible trace element patterns show depletions of Nb, Sr, and Sm, and enrichments of Th, La,
Crustal growth and deformational processes in the northern Gondwana margin
349
8º 15’
N
EMMT
+
EHMT
+
+ +
3 Km
Arraiolos
+
+ +
74
+
+
+
+
+
+
+
MNSZ
+
+
+
Cabrela
+
+ + + ARL
+ +
+
V
+
+
+
+ +
+ BDV +
+
+
Montemor-o-Novo +
CAB MAR
+
+
38º 40’
+
+
+
AMS
SIL
+
+ +
+ + +
+
REL +
+
+ + + PFAA1
CN
PTFZ
80 SEC CSN
SCD
+ + 80
CTB +
VAL
+ +
+
+ +
+ +
+
+
+ + PFMAS + + GPAS + + + + + PFPAS + + + + + + + XC EVR + + Évora MTN Valverde +
BSC MOF SEC
+
+ +
+
+
+
+
+
+
+
+ +
+ +
+
+ +
+
+ + +
+ +
+
+ +
+ +
+ +
+
+
+
+
Figure 8. Simplified structural sketch map of the Évora Massif with locations of samples for whole-rock geochemistry. EHMT—Évora highgrade metamorphic terrains; EMMT—Évora medium-grade metamorphic terrains; MNSZ—Montemor-o-Novo shear zone; PTFZ—Porto-Tomar fault zone. Other abbreviations are sample names.
Ce, Nd, and Zr. The enrichment of LREE and a pronounced characteristic negative Eu anomaly suggest a continental crust with a felsic magmatic nature and recycled detritus as the main source for these rocks. The REE abundance patterns normalized with respect to post-Archean average Australian Shale (PAAS) show flatter distribution with a positive Eu anomaly that indicates detrital plagioclase. Trace element discrimination diagram Th-Sc-Zr/10 (Bhatia and Crook, 1986; Fig. 9B) indicates that the Ediacaran detrital rock geochemical signature essentially reflects the nature of an inherited continental magmatic arc source (La/ Th = 3.9; Th/Sc = 0.67; La/Sc = 2.4, average values). This scenario began with the deposition of detrital sediments, which were predominantly of felsic provenance (TiO2 = 0.8 wt% and Ni = 45.3 ppm, average values). The Lower-Middle Cambrian micaschist (sample VAL-23) and the Upper Cambrian–Lower Ordovician (MTN-1, REL10, XC-20, BDV) have chondrite-normalized patterns similar to those of the Ediacaran Serie Negra sediments (LaN/YbN = 11.88–12.48), the fine-grained sediments having higher values that range between 18.55 and 20.94 for LaN/YbN (Fig. 9A). The TiO2 (0.64–1.19 wt%) and Ni (≤40 ppm) contents encompass similar ranges in the Cadomian sediments, indicating a felsic
provenance. A significant difference concerns the higher Al2O3/ SiO2 ratios (ranging from 0.41 to 0.46) and K2O + Na2O contents (ranging from 6.7 to 8.13 wt%) of the fine-grained sediments from the igneous (basic-dominated)-sedimentary complex relative to the Série Negra (Al2O3/SiO2 = 0.21–0.28 and K2O + Na2O = 0.63–0.97 wt%). This difference is mainly related to the significant presence of muscovite (fine-grained sediments) and muscovite + K-feldspar (coarse-grained sediments) in the Upper Cambrian–Lower Ordovician sediments. In the discrimination diagram Th-Sc-Zr/10, these rocks are plotted close to the continental arc, in transition between active continental margin and passive margin (Fig. 9B). Lower Paleozoic Magmatism: From Cadomian Accretion to a Widespread Rifting Setting The magmatic activity in the Évora Massif during the Lower Paleozoic is recorded in igneous-sedimentary complexes, where a compositional range from rhyodacites-rhyolites to subalkaline basalts is present. Volcanism started in the Early Cambrian with acidic to intermediate flows and pyroclasts, basic sills, and dikes associated with detrital and carbonate rocks. Toward the top,
350
Pereira et al. TABLE 5. REPRESENTATIVE ELECTRON MICROPROBE ANALYSIS AND STRUCTURAL FORMULAS OF FELDSPAR FROM BSC-2, SCD, REL, ARL-1, AND BDV-2
Sample Analysis (wt%) SiO2 TiO2 Al2O3 Cr2O3 FeO MnO MgO CaO Na2O K2O NiO Total (ppm) Si Al Ti Cr ++ Fe Mn Mg Ni Ca Na K XAn XAb XOr
BSC-2 294
BSC-2 297
BSC-2 299
BSC-2 302
BSC-2 303
BSC-2 306
SCD 326
SCD 329
SCD 336
SCD 339
61.75 0.00 25.08 0.00 0.62 0.03 0.12 0.80 6.25 5.38 0.00 100.022
65.66 0.00 22.45 0.00 0.24 0.06 0.04 0.71 9.51 1.83 0.00 100.496
51.91 0.00 32.69 0.00 0.82 0.07 0.11 0.64 2.56 8.63 0.00 97.43
60.40 0.00 25.47 0.05 0.00 0.09 0.00 7.13 7.73 0.06 0.00 100.923
62.40 0.03 23.65 0.02 0.12 0.00 0.01 5.35 8.63 0.07 0.00 100.276
67.69 0.01 19.82 0.07 0.12 0.00 0.00 4.39 7.54 0.08 0.02 99.743
58.71 0.02 25.52 0.03 0.16 0.03 0.00 7.42 7.19 0.12 0.00 99.208
59.43 0.02 25.71 0.00 0.10 0.00 0.00 7.49 7.42 0.11 0.00 100.276
59.60 0.00 25.63 0.05 0.12 0.05 0.01 7.63 7.35 0.10 0.02 100.563
58.32 0.01 26.45 0.02 0.20 0.01 0.01 8.43 6.69 0.16 0.11 100.414
2.761 1.322 0.000 0.000 0.023 0.001 0.008 0.000 0.038 0.542 0.307 0.043 0.611 0.346
2.878 1.160 0.000 0.000 0.009 0.002 0.002 0.000 0.033 0.808 0.102 0.035 0.856 0.108
2.667 1.326 0.000 0.002 0.000 0.003 0.000 0.000 0.337 0.662 0.003 0.336 0.660 0.003
2.759 1.233 0.001 0.001 0.004 0.000 0.001 0.000 0.253 0.740 0.004 0.254 0.742 0.004
2.964 1.023 0.000 0.003 0.004 0.000 0.000 0.001 0.206 0.640 0.004 0.242 0.753 0.005
2.642 1.354 0.001 0.001 0.006 0.001 0.000 0.000 0.358 0.628 0.007 0.361 0.633 0.007
2.646 1.349 0.001 0.000 0.004 0.000 0.000 0.000 0.357 0.641 0.006 0.356 0.638 0.006
2.647 1.342 0.000 0.002 0.005 0.002 0.001 0.001 0.363 0.633 0.006 0.363 0.632 0.006
2.601 1.391 0.000 0.001 0.008 0.000 0.000 0.004 0.403 0.579 0.009 0.407 0.584 0.009 Continued
2.426 1.801 0.000 0.000 0.032 0.003 0.008 0.000 0.032 0.232 0.515 0.041 0.298 0.661
bimodal volcanics persist, but basic occurrences increase, finally grading to massive basalts and pyroclasts that are mainly linked to detrital and minor carbonate sediments (Middle-Upper Cambrian–Lower Ordovician?). The analyses of the sampled (meta)basic rocks from both Évora medium-grade metamorphic terrains and Montemor-oNovo shear zone allowed a subdivision in three main groups: Group I displays geochemical signatures similar to those found in magmatic arc basalts, Group II shows compositional patterns typical of enriched MORB (E-MORB)-type basalts, and Group III includes metabasites derived from protoliths with normal MORB (N-MORB) compositions. The samples of Group I were collected in outcrops of the amphibolites of Casas Novas (CSN-A24), Valverde (VAL-2), and Évora (EVR-4; see Table 1A), which are intercalated in the lithologies from the Ediacaran and/or the Early Cambrian. The sampled metabasites seem to derive from basaltic—sometimes close to andesitic—compositions (Fig. 10A) and are locally associated to metarhyodacites and metasediments of both carbonate and detrital origins. The garnet-rich amphibolites from Biscaia (BSC-2, BSC-3, and BSC-10), also included in this group, crop
out within the Serie Negra metasediments. All these basic rocks show trace element patterns with pronounced negative Ta anomalies, common in basic magmas generated in the mantle wedges above subduction zones. The Th/Yb ratios are variable (from 0.39 to 0.07), corresponding to chondrite-normalized patterns that are slightly depleted, flat, or slightly enriched (Fig. 10B), which suggests a change in the trace element signatures from tholeiitic to calc-alkaline compositions. The use of the Hf/3-Th-Ta discrimination diagram (Wood, 1980) corroborates this conclusion (Fig. 10E). In the Th/Yb versus Nb/Yb diagram (Fig. 10F), this geochemical group defines a trend toward the area above the mantle array, which is also characteristic of magmatism related to subduction. The other two groups (Groups II and III; Table 1A, B and Fig. 10C, D) comprise rocks derived from essentially subalkaline basaltic protoliths (Fig. 10A), whose most striking features have been previously described by Pereira et al. (2004). First, the composition of these samples match very well with those of magmas generated in anorogenic settings characterized by low values (between 1.0 and 2.0) of the Th/Ta ratio (Fig. 10E). Second, there is a general positive correlation between Nb/Yb and Th/Yb,
Crustal growth and deformational processes in the northern Gondwana margin
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TABLE 5. REPRESENTATIVE ELECTRON MICROPROBE ANALYSIS AND STRUCTURAL FORMULAS OF FELDSPAR FROM BSC-2, SCD, REL, ARL-1, AND BDV-2 (continued) BDV-2 BDV-2 BDV-2 BDV-2 ARL-1 ARL-1 ARL-1 ARL-1 ARL-1 REL REL 1 2 3 4 1 2 3 4 5 1 2
Sample Analysis (wt%) SiO2 62.7 61.63 62.17 63.81 58.13 TiO2 0.03 0.0 0.0 0.03 0.01 23.6 23.88 24.04 22.36 26.21 Al2O3 Cr2O3 0.0 0.0 0.0 0.0 0.0 FeO 0.16 0.13 0.05 0.12 0.17 MnO 0.0 0.0 0.0 0.0 0.0 MgO 0.0 0.0 0.0 0.0 0.0 CaO 5.17 5.47 5.46 3.68 8.68 Na2O 8.13 7.98 7.99 9.06 6.33 K2O 0.05 0.02 0.03 0.05 0.04 NiO 0.0 0.0 0.0 0.0 0.0 Total 99.84 99.15 99.77 99.22 99.58 (ppm) Si 2.775 2.751 2.755 2.835 2.609 Al 1.231 1.257 1.256 1.171 1.387 Ti 0.001 0.000 0.000 0.001 0.000 Cr 0.000 0.000 0.000 0.000 0.000 ++ Fe 0.006 0.005 0.002 0.004 0.006 Mn 0.000 0.000 0.000 0.000 0.000 Mg 0.000 0.000 0.000 0.000 0.000 Ni 0.000 0.000 0.000 0.000 0.000 Ca 0.245 0.262 0.259 0.175 0.417 Na 0.698 0.691 0.687 0.780 0.551 K 0.003 0.001 0.002 0.003 0.002 XAn 0.26 0.27 0.27 0.18 0.43 XAb 0.74 0.72 0.72 0.81 0.57 XOr 0.00 0.00 0.00 0.00 0.00 Notes: XAb—albite; XAn—anorthite; XOr—orthoclase.
57.54 0.0 26.55 0.0 0.14 0.0 0.0 9.17 5.91 0.06 0.0 99.43 2.590 1.409 0.000 0.000 0.005 0.000 0.000 0.000 0.442 0.516 0.003 0.46 0.54 0.00
58.5 0.0 26.12 0.0 0.14 0.0 0.02 8.39 6.22 0.05 0.0 99.44 2.623 1.381 0.000 0.000 0.005 0.000 0.001 0.000 0.403 0.541 0.003 0.43 0.57 0.00
58.42 0.02 25.96 0.0 0.19 0.0 0.0 8.5 6.29 0.06 0.0 99.44
57.98 0.04 26.71 0.0 0.06 0.0 0.0 8.91 6.21 0.04 0.0 100.01
2.623 1.374 0.001 0.000 0.007 0.000 0.000 0.000 0.409 0.548 0.003 0.43 0.57 0.00
2.593 1.408 0.001 0.000 0.002 0.000 0.000 0.000 0.427 0.538 0.002 0.44 0.56 0.00
57.12 0. 26.43 0.0 0.25 0.0 0.01 9.18 5.61 0.05 0.0 98.65 2.589 1.412 0.000 0.000 0.009 0.000 0.001 0.000 0.446 0.493 0.003 0.47 0.52 0.00
Th 400
Sample / Chondrite
A
B
100
Active continental margin Passive margin Continental arc 10
1 La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
Sc
Zr/10
Ediacaran and Early-Middle Cambrian metasediments: CN-1, CN-2, CN-3, BSC-1, VAL-22 Upper Cambrian - Lower Ordovician metapelites: REL-10, BDV Upper Cambrian - Lower Ordovician metapsammites: MTN, XC-20 Figure 9. (A) Chondrite-normalized REE patterns for the Ediacaran to Lower Ordovician(?) sediments of the Évora Massif. Normalization values from Nakamura (1974). (B) Th-Sc-Zr/10 discrimination diagram after Bhatia and Crook (1986).
57.02 0. 26.68 0.0 0.3 0.0 0.06 9.45 5.47 0.03 0.0 99.01 2.577 1.422 0.000 0.000 0.011 0.000 0.004 0.000 0.458 0.479 0.002 0.49 0.51 0.00
REL 6 57.77 0.03 26.34 0.0 0.07 0.0 0.0 9.1 5.71 0.05 0.0 99.07 2.603 1.399 0.001 0.000 0.003 0.000 0.000 0.000 0.439 0.499 0.003 0.47 0.53 0.00
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Rock / Chondrite
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Group II (E-MORB basalts): ARL-1, ARL-2, ARL-3, ARL-10, BDV-1, BDV-2/R, REL Group III (N-MORB basalts): BSC-6, CSN-B24, CTB, MAR, MOF, SLV, VAL-1
10
D 1 Th
La Ta
Pr Ce
Sm Hf Gd Dy Ho Tm Lu Nd Zr Eu Tb Y Er Yb
Figure 10. The Évora Massif basic rocks. (A) Zr/TiO2-Nb/Y diagram of Winchester and Floyd (1977) for rocks classification. Alk-Bas—alkaline basalt; And—andesite; Bsn/Nph—basanite-phonotephrite; Com/Pant—comenditic/pantelleritic; Tachy—tachyte. (B,C,D) Multi-element patterns of selected basic rocks normalized to the chondrite of Sun and McDonough (1989), from Groups I, II, and III, respectively. (E) Plot of the studied basic rocks in the Hf/3-Th-Ta diagram of Wood (1980) and comparison with the calc-alkaline basaltic andesites from Safira. Safira eclogites described by Pedro (1996) and Leal (2001). (F) Th/Yb vs. Ta/Yb logarithmic plots with the fields of N-MORB, E-MORB, PM (primitive mantle), and OIB (oceanic island basalts). After Pearce (1983) and Sun and McDonough (1989).
Crustal growth and deformational processes in the northern Gondwana margin and all the samples (except VAL-1, which seems to represent a cumulate, as suggested by its coarse-grained texture and high-Cr content) plot within the mantle array in the diagram proposed by Pouclet et al. (1995; See Fig. 10F). Nevertheless, some differences can be pointed out in terms of the representation of the two groups on multi-elements diagrams: Group II samples have relatively enriched signatures (Th/Yb = 0.19–0.54; Ce/Yb = 4.98– 11.55), with maximum normalized values usually corresponding to the Ta position; in contrast, Group III samples display trace element patterns with large-ion lithophile element (LILE) depletion (Th/Yb = 0.05–0.14) and relatively flat normalized REE distribution (corresponding to Ce/Yb = 3.10–4.71). Therefore, whereas Group III samples are very similar to N-MORB basalts, the Group II metabasites are geochemically transitional between E-MORB and intraplate mafic volcanics, as shown in the Hf/3Th-Ta diagram (Wood, 1980). The spatial distribution of the subalkaline amphibolites from the Middle-Upper Cambrian–Lower Ordovician(?) shows some variations: the Group II, E-MORB-type, tholeiites are dominant in the Évora medium-grade metamorphic terrains, whereas compositions of the amphibolites vary between N-MORB and E-MORB types within the Montemor-o-Novo shear zone. In the Montemor-o-Novo shear zone and the Évora medium-grade metamorphic terrains, sills or dikes of N-MORB- and E-MORB-type basalts show intrusive relations with the Cadomian basement and the Early Cambrian detrital and carbonate rocks. N-MORB amphibolites were also identified as enclaves in the Évora highgrade metamorphic terrains migmatites. Variscan Synorogenic and Late Orogenic Magmatism Andesites, dacites, rhyodacites, and rhyolites occur interbedded in the Cabrela graywacke and pelite succession from a thick volcanic-sedimentary sequence with Viséan fauna (e.g., Pereira and Oliveira, 2001). Andesitic dikes (Almansor andesitic dike and Casas Novas dacitic dike) can also be found intruding the footwall Évora high-grade metamorphic terrains migmatites and the gneisses from the boundary of the Montemor-o-Novo shear zone and the Évora high-grade metamorphic terrains. The immobile element compositions of these rocks (Zr/Nb = 17.3–21.3; Zr/Y = 5.12–6.5) are analogous to the intermediate volcanics interbedded in same-age sediments from the Toca da Moura and Corte Pereiro outcrops (see Fig. 1 for location) studied in detail by Santos et al. (1987), who interpreted the genesis of these outcrops as involving a volcanic arc setting. In the Évora high-grade metamorphic terrains, tonalites, gabbros, and diorites exposed with frontiers following the northwestsoutheast strike of the Variscan structures are strictly related to local melting in the migmatitic host rocks. Amphibole and biotite separates from two samples of the Hospitais tonalites, collected near Montemor-o-Novo, yielded Ar/Ar and K/Ar ages of ca. 320 Ma. This age was interpreted as the igneous crystallization (Moita et al., 2005c) and a younger age of ca. 307 Ma obtained by K/Ar was related to the resetting caused by thermal effects of
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late orogenic granodiorite and granites. Major and trace element geochemistry reveal that this igneous association is intermediate between low-K tholeiitic and calc-alkaline series (Fig. 11A–D). The multi-element diagrams for the tonalites and gabbros are quite similar, showing moderate LILE/high field strength element (HFSE) enrichment, with ThN/YN = 2.71–11.43 for the tonalites and ThN/YN = 2.50–6.61 for the gabbros (Fig. 11B, C). These data also suggest that gabbros, diorites, and tonalites belong to a magma suite dominantly derived by fractional crystallization. Furthermore, noncumulate lithologies belonging to this sequence have major and trace element fingerprints (pronounced anomalies of Ti and Nb, with ThN/NbN = 1.33–6.22 for the tonalites and ThN/ NbN = 1.45–3.77 for the gabbros) typical of magmas generated in a suprasubduction setting (Moita et al., 2005b,c). The intrusions of foliated tonalites and granodiorites in the Évora medium-grade metamorphic terrains include microdiorite enclaves as well as previously deformed and metamorphosed paragneisses, metagraywackes, marbles, and amphibolite enclaves. Weakly foliated granitoid dikes intruding gneisses (Casa Novas granite dike) and amphibolites (Arraiolos granodiorite dike) were sampled within the Montemor-o-Novo shear zone/ Évora high-grade metamorphic terrains limit and in the Évora medium-grade metamorphic terrains, respectively. They represent granites and granodiorites with REE patterns similar to those obtained for tonalites from the Évora high-grade metamorphic terrains (Fig. 11B). The voluminous late orogenic intrusions from nearby Évora (Évora granodiorites: samples PFPASS, PFAA1, PFMAS, and granites: GPAS; see Table 1A, B and Fig. 8 for sample locations) exhibit different patterns among them (Fig. 11C). On multi-element diagrams, the granodiorites show a calc-alkaline fingerprint similar to the Hospitais tonalite, having Nb and Ti negative anomalies, whereas the peraluminous granite intimately associated with these granodiorites points to a more crustal origin (Ribeiro, 2006). The Évora granodiorites show appreciable LILE/HFSE enrichment (ThN/YN = 73.94–95.15), in contrast to a general impoverishment of REE and moderate LILE/HFSE enrichment (ThN/YN = 7.99) for the granites (Fig. 11C). Tectonic settings geochemically described in terms of HFSE and REE display a calc-alkaline signature for the intermediate dikes from Almansor and Casas Novas as part of a trend exhibiting a continuous compositional range for same-age basic flows of Toca da Moura and the acidic-intermediate differentiates from Cabrela (Fig. 11D). The Rb-Y+Nb discriminant diagram for all these sampled rocks indicates a volcanic arc signature, with the exception of the Évora granite plotted along the limit with the syncollisional field (Fig. 11E). DISCUSSION AND CONCLUSIONS The Inherited Calc-Alkaline Cadomian Arc Signature The geochemical characteristics of the Évora Massif sediments show similarities between the Neoproterozoic and the
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Almansor and Casas Novas intermediate dykes: AMS-202, CN-01 Rhyolite Com/Pan
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Cabrela acid-intermediate sills and flows: CAB-19B, CAB-23, CAB-25A
Rhyodacite-Dacite
SiO2
Trachyte
60
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Casas Novas and Arraiolos granitoid dykes: ARL-6, CN-02
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Évora porphyritic granodiorite Évora granite
Figure 11. The Évora Massif intermediate-acid rocks. (A) Zr/TiO2-SiO2 diagram of Winchester and Floyd (1977) for rocks classification. (B, C) Multi-element patterns of selected intermediate and acid rocks (synorogenic and late orogenic magmatism, respectively) normalized to the primitive mantle of Sun and McDonough (1989) and comparison with the calc-alkaline suite of gabbros and tonalites from the Évora high-grade metamorphic terrains (Hospitais Massif described by Moita et al., 2005a). (D) Plot of the studied basic rocks in the Ti/100-Zr-Y diagram of Pearce and Cann (1973) and comparison with the Viséan age calc-alkaline basaltic andesites (Toca da Moura mafic flows described by Santos et al., 1987). (E) Rb vs. Y + Nb discrimination diagram for granites of Pearce et al. (1984), showing the fields of volcanic arc granites, oceanic-ridge granites, within-plate granites, and syn-collisional granites, and comparison with calc-alkaline intermediate rocks (Corte Pereiro diorites and Hospitais tonalities, respectively described by Santos et al., 1987, and Moita et al., 2005a,b). EHMT—Évora high-grade metamorphic terrains.
Crustal growth and deformational processes in the northern Gondwana margin Lower Paleozoic record. The similarities indicate that the source for these rocks has not changed significantly with time. These sediments were deposited in a continental margin setting and derived from a dismantled Cadomian magmatic arc. Recent studies based on detrital sediment geochemistry show the existence of a strong inherited calc-alkaline arc affinity, derived from the development of an active plate margin, which was preserved on such immature synorogenic Neoproterozoic and even synorogenic and/or late-orogenic Early Cambrian sediments (e.g., Dabard, 1990; Linnemann and Romer, 2002; Pereira and Chichorro, 2004; Pereira et al., 2006). In the Évora Massif, the Ossa-Morena zone Neoproterozoic (Serie Negra) crops out with several characteristics that can be correlated with the Cadomian geodynamic evolution fingerprints of contemporaneous segments of western and central Europe (North Armorican Cadomian belt, Saxo-Thuringian zone), Morocco (Western Meseta, High- and Anti Atlas), and others from the west and east Avalonian parts of such old orogenic belts (southeastern New England, southern New Brunswick, Nova Scotia/Cape Breton Island, eastern Newfoundland, north Wales and southeastern Ireland, and the British midlands) (e.g., Eguiluz et al., 2000; Linnemann et al., 2000, 2004; Murphy et al., 2002; Nance et al., 2002; Pereira et al., 2003b, 2006). These authors consider that the late Neoproterozoic–Early Cambrian evolution records a transition from an active plate margin, with intense and heterogeneous arc magmatism and possible events of arc-related rifting, to an intracontinental transcurrent setting often associated with more bimodal magmatism. Early Paleozoic Crustal Thinning, Fragmentation, and Anorogenic Magmatism Field and geochemical data from Lower Paleozoic basic magmatic rocks of the Évora Massif allowed subdivision of these rocks into three main groups with two distinct origins: (1) Group I, made up of metabasites with geochemical signatures typical of convergent plate margins formed during late stages of the Cadomian orogeny or the initial stages of the Cambrian rifting; and (2) Groups II and III, with tholeiitic MORB-like composition metabasites suggesting a significant rifting period during the Upper Cambrian–Lower Ordovician(?). To explain the more enriched source, the Nb and Ta negative anomaly, and the strong increase in the Th/Yb ratio in the samples of basic rocks from Group I, two types of hypothesis are plausible: a “subduction component” or a mixture between MORB basalts and crustal melts (Wilson, 1989). Until now rocks with similar composition have been interpreted as continental tholeiites originated from a contaminated magma with an enriched-MORB source and linked to a process of assimilation-fractional crystallization on a calc-alkaline continental crust that suffered rifting (e.g., Safira eclogites; Leal et al., 1997; Leal, 2001). We interpret the formation of these basic rocks as linked to petrogenetic processes involving addition of a calc-alkaline magma, which was extracted from a depleted mantle reservoir
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related with a continental arc. In fact, they occur associated with felsic volcanics of calc-alkaline tendency in the Early Cambrian carbonate and detrital sequence or in direct relation with the Ediacaran sediments (Serie Negra). The association is due to mixed processes that are probably related with the latest stages of the Cadomian orogeny. Groups II and III, with MORB-like compositions, represent a basic magmatism closely associated with a variable-thickness monotonous pelitic series that has important pyroclastic input. These igneous-sedimentary complexes are typical of a continental margin setting affected by strong stretching and evolving to an active rift linked to asthenosphere uplift and intense magmatism. These amphibolites may indicate the development of spreading-center depleted sources (N-MORB) contemporaneous with plume-dominated enriched sources (E-MORB) (Pedro, 2004; Pereira et al., 2004). Transtension with heterogeneous spatial and temporal distribution of stretching was probably responsible for the introduction of great variations in the crust fragmentation and thickness, creating a favorable geodynamic scenario for a wide range of degrees of contamination. The proposed stratigraphy for the Cambrian of the Èvora Massif has similarities with the sequences from the Zafra and Jerez de los Caballeros sectors in Spain (Extebarria et al., 2006). Features of the stratigraphy and geochemistry for the igneoussedimentary complexes testify to an important crustal extension event probably related to the opening of the Rheic Ocean (e.g., Sanchez Garcia et al., 2003). Upper Paleozoic Migmatization, Transcurrent Tectonics, and Orogenic Magmatism Evidence for the transition from overall crustal shortening to extensional tectonism in the northern Gondwana margin is locally preserved in the Évora Massif. Here, high- to medium-pressure metamorphic rocks, which were formed during a collision process at the northern Gondwana margin, have been exhumed along high-strain ductile shear zones (probably low-angle detachments with orogen-parallel or slightly oblique movement). The Évora Massif tectonites present a strong mineral stretching lineation that trends generally N330°–340° to N300°–310°. The associated foliation is affected on the kilometer scale by open upright late folds. Shear-sense criteria (C-S structures, extensional shear bands, and asymmetric pressure shadows) along the mineral stretching lineation record a dominant sinistral shearing with top-toward-the-northwest, when the foliation deeps to northeast, or top-to-southeast with the foliation dipping to the southwest. They also record local dextral shearing or symmetric fabrics. Mesoscale folds show axes more or less parallel to the stretching lineation. These structures are related to a progressive and complex shearing that accompanied the generalized amphibolite-facies metamorphism. Evidence for metamorphism with an increase in pressure while maintaining a moderate temperature exists in the zoning of garnets in the Biscaia amphibolites and in the inclusions of
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garnets in the Safira eclogites. The eclogites, which crop out in meter-scale pods, are wrapped around by more deformed and partially retrogressed rocks with mineral assemblages, which crystallized and/or recrystallized to form S-L fabrics. Shear-sense indicators can be related to extension and exhumation. Toward the northeast, metamorphism, with increasing temperature and isobaric low-pressure conditions, is dominant. Here the steeply dipping (Casas Novas-Boa Fé) and the flat-lying (Évora) foliations and horizontally disposed stretching lineations—as well as the development of mesoscale folds having axes parallel to the stretching lineation and to the long-axes of boudins—indicate transtension during uplift. Metamorphism and shearing show a close spatial and temporal relationship with migmatization and subsequent magmatism as the plausible result of a continuous process of heat advection. This advection resulted from the flow of melt sources directly from mantle-derived basic rocks or from a subduction slab (with sediments and oceanic basalts) and the overlying mantle lithosphere. Thus, the Variscan evolution of the Évora Massif crust during the Upper Paleozoic could not have evolved by a simple switch from early crustal thickening to later overall thinning; instead, it must have undergone multiple oscillations between shortening and crustal extension (due to transcurrent movements), causing important perturbations in the thermal framework. Transcurrent movements involve crustal extension and formation of major ductile shear zones, resulting in the partial exhumation of the Évora high-grade metamorphic terrains. A consequence of this process is the huge volume of melt produced by uplift and decompression, as evidenced by the widespread orogenic magmatism intruded on the gneisses and migmatites (footwall) of the Évora high-grade metamorphic terrains, which occurred after a period of quiescence of migmatization and the intrusion of tonalites and gabbros (ca. 320 Ma). This early intermediate-basic magmatism first intruded a previous hot crust, causing local partial melting. Finally, a sufficient volume of magma was accumulated to cause its emplacement along the boundary to the Évora high-grade and medium-grade metamorphic terrains (major detachment) and, at shallow crustal levels, the Évora medium-grade metamorphic terrains (hanging wall). Based on regional comparisons, the main regional structures observed within the Ossa-Morena zone are related to strong crustal extension and shearing. The recognition of fabrics related to the Variscan plate convergence and the Lower Paleozoic basin inversions is still controversial. To explain these structures, some authors believe that the widespread extensional fabrics were largely controlled by the structures inherited from crustal thickening (e.g., Giese et al., 1994) and were responsible for erasing the former structures (e.g., Díaz Azpiroz et al., 2004). The extension is connected with basin deposition and calc-alkaline volcanism and/or plutonism was concomitant with continuing partial exhumation of high-grade metamorphic terrains during the Lower Carboniferous. Similar-age migmatization and deformation processes are observed in the Aracena Massif (Spain; Castro et al., 1999; Díaz Azpiroz et al., 2004).
ACKNOWLEDGMENTS This work was supported by the Geophysical Centre of Évora (University of Évora)/Portuguese Science and Technology Foundation (FCT) and funds from research projects CGL2004– 06808-CO4/BTE (Universities of Huelva, Pablo de Olavide, and Évora) and MODELIB 35630/99 (University of Aveiro). This work is a contribution to the International Geological Correlation Program Projects 497 and 485. M. Chichorro benefits from a FCT/PRAXXIS research doctoral scholarship. REFERENCES CITED Bhatia, M.R., and Crook, K.A.W., 1986, Trace element characteristics of greywackes and tectonic setting discrimination of sedimentary basins: Contributions to Mineralogy and Petrology, v. 92, p. 181–193, doi: 10.1007/ BF00375292. Blatt, H., Middleton, G., and Murray, R., 1980, Origin of sedimentary rocks: New York, Prentice Hall, 634 p. Brown, E.H., 1974, Comparison of the mineralogy and phase relations of blueschists from the northern Cascades, Washington, and the blueschists from Otago, New Zealand: Geological Society of America Bulletin, v. 85, p. 333–344. Cartier, C., Faure, M., and Lardeux, H., 2002, Discussion on “A new model for the Hercynian orogen of Gondwanan France and Iberia” by D. Shelley & G. Bossière, Journal of Structural Geology 2000, v. 22, 757–776: Journal of Structural Geology, v. 24, p. 1659–1663, doi: 10.1016/S01918141(01)00150-X. Carvalhosa, A., 1983, Esquema geológico do Maciço de Évora: Comunicações dos Serviços Geológicos de Portugal, v. 69, no. 2, p. 201–208. Carvalhosa, A., 1999, Carta Geológica de Portugal, Noticia Explicativa da Folha 36-C (Arraiolos): Lisbon, Instituto Geológico e Mineiro, scale 1:50,000. Carvalhosa, A., and Zbyszewski, G. 1994, Carta Geológica de Portugal, Noticia Explicativa da Folha 35-D (Montemor-o-Novo): Lisbon, Instituto Geológico e Mineiro, scale 1:50,000. Carvalhosa, A., Galopim de Carvalho, A.M., Matos Alves, C.A., and Pina, H.L., 1969, Carta Geológica de Portugal, Noticia Explicativa da Folha 40-A (Évora): Lisbon, Serviços Geológicos de Portugal, scale 1:50,000. Carvalhosa, A., Gonçalves, F., and Oliveira, V., 1987, Carta Geológica de Portugal, Noticia Explicativa da Folha 36-D (Redondo): Lisbon, Instituto Geológico e Mineiro, scale 1:50,000. Castro, A., Fernández, C., El-Hmidi, H., El-Biad, M., Díaz, M., de la Rosa, J., and Stuart, F., 1999, Age constraints to the relationship between magmatism, metamorphism and tectonism in the Aracena metamorphic belt, southern Spain: International Journal of Earth Sciences, v. 88, p. 26–37, doi: 10.1007/s005310050243. Chichorro, M., 2006, Estrutura do Sudoeste da Zona de Ossa-Morena: Área de Santiago de Escoural–Cabrela (Zona de Cisalhamento de Montemoro-Novo, Maciço de Évora) [Ph.D. diss.]: Évora, Portugal, Universidade de Évora. Chichorro, M., Pereira, M.F., Apraiz, A., and Silva, J.B., 2003, Syntectonic high-temperature/low-pressure metamorphism in the Boa Fé fault zone (Évora Massif, western Ossa-Morena zone), in Proceedings VI Congresso Nacional de Geologia, Tema D—Tectónica, Geofísica, Monte da Caparica, 4–6 June: Lisbon, Universidade Nova Lisboa, p. D18–D20. Chichorro, M., Pereira, M.F., Apraiz, A., and Silva, J.B., 2004, High temperature–low pressure tectonites from the Boa Fé fault zone (Évora Massif, Ossa-Morena zone, Portugal): Evidence for transtensional tectonics: Geogaceta, v. 34, p. 43–46. Dabard, M.P., 1990, Lower Brioverian formations (Upper Proterozoic) of the Armorican Massif (France): Geodynamic evolution of source areas revealed by sandstone petrography and geochemistry: Sedimentary Geology, v. 69, p. 45–58, doi: 10.1016/0037-0738(90)90100-8. Dewey, J.F., Holdsworth, R.E., and Strachan, R.A., 1998, Transpression and transtension zones, in Holdsworth, R.E., Strachan, R.A., and Dewey, J.F., eds., Continental transpressional and transtensional tectonics: London, Geological Society of London Special Publication 135, p. 1–14.
Crustal growth and deformational processes in the northern Gondwana margin Dias, R., and Ribeiro, A., 1995, The Ibero-Armorican arc: A collision effect against an irregular continent? Tectonophysics, v. 246, p. 113–128, doi: 10.1016/0040-1951(94)00253-6. Díaz Azpiroz, M.D., Castro, A., Fernandez, C., López, S., Fernandez Caliani, J.C., and Moreno-Ventas, I., 2004, The contact between the Ossa Morena and the South Portuguese zones. Characteristics and significance of the Aracena metamorphic belt, in its central sector between Aroche and Aracena (Huelva): Journal of Iberian Geology, v. 30, p. 23–51. Diáz Garcia, F., Martínez Catalán, J.R., Arenas, R., and González Cuadra, P., 1999, Structural and kinematic analysis of the Corredoiras detachment: Evidence for early Variscan synconvergent extension in the Órdenes complex, NW Spain: International Journal of Earth Sciences, v. 88, p. 337– 351, doi: 10.1007/s005310050269. Dörr, W., Zulauf, G., Fiala, J., Franke, W., and Vejnar, Z., 2002, Neoproterozoic to Early Cambrian history of an active plate margin in the Teplá-Barrandian unit—A correlation of U-Pb isotopic-dilution-TIMS ages (Bohemia, Czech Republic): Tectonophysics, v. 352, p. 65–85, doi: 10.1016/S00401951(02)00189-0. Eguiluz, L., Gil-Ibarguchi, J.I., Abalos, B., and Apraiz, A., 2000, Superposed Hercynian and Cadomian orogenic cycles in the Ossa-Morena zone and related areas of the Iberian Massif: Geological Society of America Bulletin, v. 112, no. 9, p. 1398–1413, doi: 10.1130/0016-7606(2000)112<1398: SHACOC>2.0.CO;2. Extebarria, M., Chalot-Prat, F., Apraiz, A., and Eguiluz, L., 2006, Birth of volcanic passive margin in Cambrian time: Rift paleogeography of the OssaMorena zone: Precambrian Research, v. 147, p. 366–386. Floyd, P.A., Winchester, J.A., Seston, R., Kryza, R., and Crowley, Q.G., 2000, Review of geochemical variation in Lower Paleozoic metabasites from the NE Bohemian Massif: Intracratonic rifting and plume-ridge interaction, in Franke, W., Haak, V., Oncken, O., and Tanner, D., eds., Orogenic processes: Quantification and modelling in the Variscan belt of Central Europe: London, Geological Society of London Special Publication 179, p. 155–174. Franke, W., Haak, V., Oncken, O., and Tanner, D., eds., 2000, Orogenic processes: Quantification and modelling in the Variscan belt of Central Europe: London, Geological Society of London Special Publication 179, 459 p. Giese, U., Rainer, H., Hollman, G., and Walter, R., 1994, Geology of the southwestern Iberian Meseta I. The Paleozoic of the Ossa Morena zone north and south of the Olivenza-Monesterio anticline (Huelva Province, SW Spain): Neues Jahrbuch für Geologie und Paläontologie Abhandlungen, v. 192, p. 293–331. Gonçalves, F., and Carvalhosa, A., 1994, O Proterozoico da Zona de OssaMorena: Síntese e actualização de conhecimentos: Memorais Academia Ciências, Lisboa, v. 34, p. 1–35. Kohn, M.J., and Spear, F.S., 1990, Two new geobarometers for garnet-amphiboles, with applications to southwestern Vermont: American Mineralogy, v. 75, p. 89–96. Leake, B.E., 1978, Nomenclature of amphiboles: Canadian Mineralogy, v. 16, p. 501–520. Leal, N., 2001, Estudo petrológico e geoquímico de rochas metamórficas de alta pressão das regiões de Alvito-Viana do Alentejo e de Safira (Zona de Ossa-Morena, Maciço Ibérico) [Ph.D. diss.]: Lisbon, Universidade de Lisboa, 402 p. Leal, N., Pedro, J., Fonseca, P., Araújo, A., and Munha, J.M., 1997, Metamorfismo nos sectores meridionais da Zona de Ossa-Morena; Actualização de conhecimentos, in Araújo, A., and Pereira, M.F., eds., Estudos sobre a Geologia da Zona de Ossa Morena (Maciço Ibérico): Livro de Homenagem ao Prof. Francisco Gonçalves: Évora, Portugal, Universidade Évora, p. 119–132. Linnemann, U., and Romer, R.L., 2002, The Cadomian orogeny in SaxoThuringia, Germany: Geochemical and Nd-Sr-Pb isotopic characterization of marginal basins with constraints to geotectonic setting and provenance: Tectonophysics, v. 352, p. 33–64, doi: 10.1016/S0040-1951(02)00188-9. Linnemann, U., Gehmlich, M., Tichomirowa, M., Buschmann, B., Nasdala, L., Jonas, P., Lützer, H., and Bombach, K., 2000, From Cadomian subduction to Early Paleozoic rifting: The evolution of Saxo-Thuringia at the margin of Gondwana in the light of single zircon geochronology and basin development (central European Variscides, Germany), in Franke, W., Haak, V., Oncken, O., Tanner, D., eds., Orogenic processes: Quantification and modelling in the Variscan belt of Central Europe: London, Geological Society of London Special Publication 179, p. 131–153. Linnemann, U., McNaughton, N.J., Romer, R.L., Gehmlich, M., Drost, K., and Tonk, C., 2004, West African provenance for Saxo-Thuringia
357
(Bohemian Massif): Did Armorica ever leave pre-Pangean Gondwana?: U/Pb-SHRIMP zircon evidence and the Nd-isotopic record: International Journal of Earth Sciences, v. 93, p. 683–705, doi: 10.1007/s00531-0040413-8. Martínez Catalán, J.R., Díaz Garcia, F., Arenas, R., Abati, J., Castiñeiras, P., Gonzalez Cuadra, P., Gómez Barreiro, J., and Rubio Pascual, F.J., 2002, Thrusts and detachments systems in the Órdones complex (northwestern Spain): Implications for the Variscan-Appalachian geodynamics, in Martínez Catalán, J.R., Hatcher, R.D., Arenas, R., and Díaz Garcia, F., eds., Variscan-Appalachian dynamics: The building of the late Paleozoic basement: Boulder, Colorado, Geological Society of America Special Paper 364, p. 57–71. Matte, Ph., 1991. Accretionary history and crustal evolution of the Variscan belt in Western Europe, in Hatcher, R.D., Jr., and Zonenshain, L., eds., Accretionary tectonics and composite continents: Tectonophysics, v. 196, p. 309–337. Matte, Ph., 2001, The Variscan collage and orogeny (480–290 Ma) and the tectonic definition of the Armorica microplate: A review: Terra Nova, v. 13, p. 122–128, doi: 10.1046/j.1365-3121.2001.00327.x. Moita, P., Pereira, M.F., and Santos, J.F., 2005a, Tonalites from the Hospitais Massif (Ossa-Morena zone, SW Iberian Massif, Portugal). I: Geological setting and petrography: Geogaceta, v. 37, p. 51–54. Moita, P., Santos, J.F., and Pereira, M.F., 2005b, Tonalites from the Hospitais Massif (Ossa-Morena Zone, SW Iberian Massif, Portugal). II: Geochemistry and petrogenesis: Geogaceta, v. 37, p. 55–58. Moita, P., Santos, J.F., and Pereira, M.F., 2005c, Dados geocronológicos de rochas intrusivas sin-tectónicas do maciço dos Hospitais (Montemor-oNovo, Zona de Ossa-Morena), in Proceedings VIII Congresso de Geoquímica de Países de Língua Portuguesa, vol. 2: Aveiro, Portugal, Universidade Aveiro, p. 471–474. Moores, E.M., and Twiss, R.J., 1995, Tectonics: New York, W.H. Freeman & Co., 415 p. Murphy, J.B., Eguiluz, L., and Zulauf, G., 2002, Cadomian orogens, peri-Gondwanan correlatives and Laurentia-Baltica connections: Tectonophysics, v. 352, p. 1–9, doi: 10.1016/S0040-1951(02)00186-5. Nakamura, N., 1974, Determination of REE, Ba, Fe, Mg, Na and K in carbonaceous and ordinary chondrites: Geochemical and Cosmochimical Acta, v. 38, p. 757–775, doi: 10.1016/0016-7037(74)90149-5. Nance, R.D., Murphy, J.B., and Keppie, J.D., 2002, A cordilleran model for the evolution of Avalonia: Tectonophysics, v. 352, p. 11–31, doi: 10.1016/ S0040-1951(02)00187-7. Oliveira, J.T., 1992, coord., Carta Geológica de Portugal: Lisbon, Instituto Geológico e Mineiro, scale 1:500,000. Oliveira, J.T., Oliveira, V., and Piçarra, J.M., 1991, Traços gerais da evolução tectonoestratigrafica da Zona de Ossa Morena em Portugal: Cuadernos Laboratorio Xeologico Laxe, Coruña, v. 16, p. 221–250. Paris, F., and Robardet, M., 1994, Paleogeographic synthesis, in Keppie, J.D., ed., Pre-Mesozoic geology in France and related areas: Berlin, SpringerVerlag, p. 172–176. Pearce, J.A., 1983, Role of the sub-continental lithosphere in magma genesis at active continental margins, in Hawkesworth, C.J., and Norry, M.J., eds., Continental basalts and mantle xenoliths: Nantwich, Shiva, p. 230–249. Pearce, J.A., and Cann, J.R., 1973, Tectonic setting of basic volcanic rocks determined using trace element analysis: Earth and Planetary Sciences Letters, v. 19, p. 290–300, doi: 10.1016/0012-821X(73)90129-5. Pearce, J.A., Harris, N., and Tindle, A.G., 1984, Trace element discrimination diagrams for the tectonic interpretation of granitic rocks: Journal of Petrology, v. 25, p. 956–983. Pedro, J., 1996, Estudo do metamorfismo de alta pressao na área de Safira (montemor-o-Novo) Zona de Ossa-Morena [M.S. diss.]: Lisbon, Universidade de Lisboa, 69 p. Pedro, J., 2004, Estudo geológico e geoquímico das Sequências Ofiolíticas Internas da Zona de Ossa-Morena (Portugal) [Ph.D. diss.]: Évora, Portugal, Universidade de Évora, 225 p. Perchuck, L.L., 1991, Derivation of a thermodynamically consistent set of geothermometers and geobarometers for metamorphic and magmatic rocks, in Perchuck, L.L., ed., Progress in metamorphic and magmatic petrology: Cambridge, Cambridge University Press, p. 93–111. Pereira, M.F., and Chichorro, M., 2004, Tectonic setting and provenance of clastic sediments from the Ossa-Morena Zone Neoproterozoic and Early Cambrian basins (Iberian Massif, Portugal, Alentejo): Geotemas, v. 6, no. 2, p. 161–164.
358
Pereira et al.
Pereira, Z., and Oliveira, J.T., 2001, The Cabrela syncline stratigraphic sequence, Ossa Morena zone, Portugal: A palynostratigraphic review, in Proceedings of the 15th International Senckenberg Conference, International Geological Correlation Program Project 421, and Subcomission on Devonian Stratigraphy, abstracts, vol. 76: Frankfurt, International Geological Correlation Program. Pereira, Z., and Oliveira, J.T., 2003, Estudo palinoestratigráfico do sinclinal da Estação de Cabrela. Implicações tectono-estratigráficas: Ciências da Terra (Universidade Nova de Lisboa), v. 5, p. A18–A19. Pereira, M.F., and Silva, J.B., 2001, A new model for the Hercynian orogen of Gondwanan France and Iberia: Discussion: Journal of Structural Geology, v. 23, p. 835–838, doi: 10.1016/S0191-8141(00)00148-6. Pereira, M.F., and Silva, J.B., 2002, The geometry and kinematics of enclaves in sheared migmatites from the Evora Massif, Ossa-Morena zone (Portugal): Geogaceta, v. 31, p. 199–202. Pereira, M.F., Silva, J.B., and Chichorro, M., 2003a, Internal structure of the Évora high-grade terrains and the Montemor-o-Novo shear zone (OssaMorena zone, Portugal): Geogaceta, v. 33, p. 79–82. Pereira, M.F., Linnemann, U., Silva, J.B., Kröner, U., and Chichorro, M., 2003b, Correlation of Neoproterozoic to Cambro-Ordovician rock units of the Ossa-Morena, the Central-Iberian and the Saxo-Thuringian zones (Portugal, Germany), in Proceedings VI Congresso Nacional de Geologia, Tema A—Paleontologia, Estratigrafia, Monte da Caparica, 4–6 June: Lisbon, Universidade Nova Lisboa, p. A114–A117. Pereira, M.F., Chichorro, M., Santos, J.F., Moita, P., and Silva, J.B., 2004, Geochemistry of lower Paleozoic anorogenic basic rocks from the Évora Massif (western Ossa-Morena zone, Portugal): Geogaceta, v. 35, p. 87–91. Pereira, M.F., Silva, J.B., and Chichorro, M., 2005, Occurrences of MORB-type magmatism in the northern Gondwana margin and their geodynamic significance: The case of the Ossa-Morena zone (SW Iberian Massif), IGCP 497: The Rheic Ocean—Its origin, evolution and correlatives, in Project meeting, conference abstracts and program, Portsmouth, 6–11 July: Portsmouth, New Hampshire, University of Portsmouth, p. 15–16. Pereira, M.F., Chichorro, M., Linnemann, U., Eguiluz, L., and Silva, J.B., 2006, Inherited arc signature in Ediacaran and Early Cambrian basins of the Ossa-Morena zone (Iberian Massif, Portugal): Paleographic link with European and north African Cadomian correlatives: Precambrian Research, v. 144, p. 297–315. Pettijohn, F.J., Potter, P.E., and Siever, R., 1972, Sand and sandstones: Berlin, Springer-Verlag, 618 p. Pouclet, A., Lee, J.S., Vidal, P., Cousens, B., and Bellon, H., 1995, Cretaceous to Cenozoic volcanism in South Korea and in the Sea of Japan: Magmatic constraints on the opening of the back-arc basin, in Smelie, J.L., ed., Volcanism associated with extension at consuming plate margins: London, Geological Society of London Special Publication 81, p. 169–191. Powell, R., 1985, Geothermometry and geobarometry: A discussion: Journal of the Geological Society of London, v. 142, p. 29–38. Quesada, C., and Munhá, J.M., 1990, Ossa-Morena zone: Metamorphism, in Dallmeyer, R.D., and Martinez Garcia, E., eds., Pre-Mesozoic geology of Iberia: Berlin and Heidelberg, Springer-Verlag, p. 314–320. Raase, P., 1974, Al and Ti contents of hornblende, indicators of pressure and temperature of regional metamorphism: Contributions to Mineral Petrology, v. 45, p. 231–236, doi: 10.1007/BF00383440. Ribeiro, A., 1983, Relações entre formações do Devónico superior e o Maciço de Évora na região de Cabrela (Vendas Novas): Comunicações dos Serviços Geológicos de Portugal, v. 69, no. 2, p. 267–282. Ribeiro, A., Quesada, C., and Dallmeyer, R.D., 1990, Geodynamic evolution of the Iberian Massif, in Dallmeyer, R.D., and Martínez García, E., eds., Pre-Mesozoic geology of Iberia: Berlin and Heidelberg, Springer-Verlag, p. 399–409. Ribeiro, C., 2006. Cartografia de rochas ígneas: Análise da orientação preferencial de fenocristais de feldspato no granitóideporfiróide do Alto de São Bento (Maciço de Évora, Zona de Ossa-Morena) [M.S. diss.]: Évora, Portugal, University of Évora, 105 p. Robardet, M., 2002, Alternative approach to the Variscan belt in southwestern Europe: Preorogenic paleobiogeographical constraints, in Martinez
Catalán, J.R., Hatcher, R.D., Jr., Arenas, R., and Diaz Garcia, F., eds., Variscan-Appalachian dynamics: The building of the late Paleozoic basement: Boulder, Colorado, Geological Society of America Special Paper 364, p. 1–15. Robardet, M., 2003, The Armorica “microplate”: Fact or fiction? Critical review of the concept and contradictory palaeobiogeographical data: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 195, p. 125–148, doi: 10.1016/S0031-0182(03)00305-5. Rollinson, H., 1993, Using geochemical data: Evaluation, presentation, interpretation: Upper Saddle River, N.J., Prentice Hall, p. 1–352. Sanchez Garcia, T., Bellindo, F., and Quesada, C., 2003, Geodynamic setting and geochemical signatures of Cambrian-Ordovician rift-related igneous rocks (Ossa-Morena zone, SW Ibéria): Tectonophysics, v. 365, p. 233–255. Santos, J.F., Mata, J., Gonçalves, F., and Munhá, J.M., 1987, Contribuição para o conhecimento geológico-petrológico da região de Santa Suzana: o Complexo Vulcano-sedimentar de Toca da Moura: Comunicações Serviços Geológicos Portugal Lisboa, v. 73, p. 29–40. Shelley, D., and Bossière, G., 2000, A new model for the Hercynian orogen of Gondwanan France and Iberia: Journal of Structural Geology, v. 22, p. 757–776, doi: 10.1016/S0191-8141(00)00007-9. Shelley, D., and Bossière, G., 2001, A new model for the Hercynian orogen of Gondwanan France and Iberia: Reply: Journal of Structural Geology, v. 23, p. 839–841, doi: 10.1016/S0191-8141(00)00147-4. Shelley, D., and Bossière, G., 2002, Reply to discussion on “A new model for the Hercynian orogen of Gondwanan France and Iberia”: Journal of Structural Geology, v. 24, p. 1665–1668, doi: 10.1016/S0191-8141(01)00151-1. Silva, J.B., and Pereira, M.F., 2004, Transcurrent continental tectonics model for the Ossa-Morena zone Neoproterozoic-Paleozoic evolution (SW Iberian Massif, Portugal): International Journal of Earth Sciences, v. 93, p. 886–896, doi: 10.1007/s00531-004-0424-5. Silva, J.B., Pereira, M.F., and Chichorro, M., 2003, Upper Paleozoic basins development under orogen-parallel sinistral transcurrent regime in the SW Iberian Massif (Portugal), in Proceedings VI Congresso Nacional de Geologia, Tema C—Bacias sedimentares, geologia estrutural, Monte da caparica, 4–6 June: Lisbon, Universidade Nova Lisboa, p. C79–C82. Spear, F.S., 1981, Amphibole-plagioclase equilibria, an empirical model for the relation albite+tremolite=edenite+4quartz: Contributions to Mineral Petrology, v. 77, p. 335–364. Stampfli, G., 2000, Tethyan oceans, in Bozkurt, E., Winchester, J.A., and Piper, J.D.A., eds., Tectonics and magmatism in Turkey and the surrounding area: London, Geological Society of London Special Publication 173, p. 1–23. Sun, S.S., and McDonough, W.F., 1989, Chemical and isotopic systematics of oceanic basalts: Implications for mantle composition and processes, in Saunders, A.D., and Norry, M.J., eds., Magmatism in ocean basins: London, Geological Society of London Special Publication 42, p. 313–345. Tait, J.A., Bachtadse, V., Franke, W., and Soffel, H.C., 1997, Geodynamic evolution of the European Variscan fold belt: Palaeomagnetic and geological constraints: Geologische Rundschau, v. 86, p. 585–589, doi: 10.1007/ s005310050165. Wilson, M., 1989, Igneous petrogenesis—A global tectonic approach: New York, Chapman & Hall, 466 p. Winchester, J.A., and Floyd, P.A., 1977, Geochemical discrimination of different magma series and their differentiation products using immobile elements: Chemical Geology, v. 20, p. 325–343, doi: 10.1016/00092541(77)90057-2. Windley, B.F., 1997, The evolving continents (3rd edition): New York, Wiley, 526 p. Wood, D.A., 1980, The application of a Th-Hf-Ta diagram to problems of tectonomagmatic classification and to establishing the nature of crustal contamination of basaltic lavas of the British Tertiary volcanic province: Earth and Planetary Sciences Letters, v. 50, p. 11–30. MANUSCRIPT ACCEPTED BY THE SOCIETY 3 OCTOBER 2006
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Geological Society of America Special Paper 423 2007
The Lower–Middle Cambrian boundary in the Mediterranean subprovince Rodolfo Gozalo* Departamento de Geología, Universitat de València. C/ Dr. Moliner 50, E-46100 Burjasot, Spain Eladio Liñán María Eugenia Dies Álvarez José Antonio Gámez Vintaned Área y Museo de Paleontología, Departamento de Ciencias de la Tierra, Universidad de Zaragoza, E-50009 Zaragoza, Spain Eduardo Mayoral Departamento de Geodinámica y Paleontología, Universidad de Huelva. Avda. de las Fuerzas Armadas s/n, E-21006 Huelva, Spain
ABSTRACT The position of the Lower–Middle Cambrian boundary, as classically used, is an issue still under discussion by the International Subcommission on Cambrian Stratigraphy (ISCS), and no level has been established yet. At present, there are two oryctocephalid trilobite species–based Global Standard Stratotype-section and Point (GSSP) proposals in the literature, Oryctocephalus indicus and Ovatoryctocara granulata. These two species have not yet been found in the Mediterranean subprovince. For this reason, other correlation tools that approximate these levels are needed. A complete chronostratigraphy for the Lower and Middle Cambrian Series of Iberia was proposed by Sdzuy (1971a,b). Recently Geyer and Landing (2004) made a new proposal for the Lower–Middle Cambrian chronostratigraphy of West Gondwana. They proposed the Agdzian Stage. This stage is more or less equivalent to the Bilbilian and Leonian Stages of Spain. The main problem with the Agdzian Stage is that its boundaries are not correlative with a potential global Cambrian Series boundary. Those levels would be in the middle part of the Agdzian Stage. In order to make a more accurate correlation, we prefer to use the previous Spanish scale and try to clarify the correlation between different sequences of this time interval in the Mediterranean subprovince. Here we present a summary of biostratigraphical data from the Bilbilian and Leonian Stages, the boundary of which is placed at the first appearance datum *E-mail:
[email protected]. Present address, Dies Álvarez: Department of Geology, GeoBiosphere Science Center, Lund University, Sölvegatan 12, SE-22362 Lund, Sweden. Gozalo, R., Liñán, E., Dies Álvarez, M.E., Gámez Vintaned, J.A., and Mayoral, E., 2007, The Lower–Middle Cambrian boundary in the Mediterranean subprovince, in Linnemann, U., Nance, R.D., Kraft, P., and Zulauf, G., eds., The evolution of the Rheic Ocean: From Avalonian-Cadomian active margin to AlleghenianVariscan collision: Geological Society of America Special Paper 423, p. 359–373, doi: 10.1130/2007.2423(17). For permission to copy, contact
[email protected]. ©2007 Geological Society of America. All rights reserved.
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Gozalo et al. (FAD) of Acadoparadoxides mureroensis. This, the classical Lower–Middle Cambrian boundary in the Paradoxides realm, is constrained by the data of various areas of the Mediterranean region. A revised correlation chart comparing the Mediterranean and other regions is presented. The position of the Bilbilian–Leonian boundary roughly coincides with the two GSSPs proposed by the ISCS. Keywords: stratigraphy, biochronology, trilobites, global correlation, Cambrian GSSP
INTRODUCTION The Cambrian is one of the systems without a complete succession of series and stages recognized globally by the International Subcommission on Cambrian Stratigraphy (ISCS) (Geyer and Shergold, 2000; Peng et al., 2004; Babcock et al., 2005). In part this is due to the provinciality of the trilobites and other faunas. Regional stages have been proposed for several areas (e.g., Siberia, Laurentia, Morocco, Australia, Baltica, China, Iberia). For Iberia, a complete chronostratigraphy was proposed for the Lower and Middle Cambrian Series (Fig. 1) by Sdzuy (1971a,b). Recently Geyer and Landing (2004) made a new proposal for the Lower–Middle Cambrian chronostratigraphy of West Gondwana; they use different stages and series previously defined in Morocco, Spain, and France. They proposed the new Agdzian Stage, which is more or less equivalent to the Bilbilian and Leonian Stages of Spain (Sdzuy, 1971a; Liñán et al., 1993a; Sdzuy et al., 1996, 1999). The main problem with the Agdzian Stage is that its boundaries do not correlate with the Oryctocephalus indicus or Ovatoryctocara granulata levels, the levels that are being considered as the base of a global Cambrian Series. Those levels would be in the middle part of the Agdzian Stage. In order to make a more accurate correlation, we prefer to use the previous Spanish scale and try to clarify the correlation between the different sequences of this time interval in the Mediterranean subprovince. The Bilbilian and Leonian Stages are chronostratigraphic units that were formerly defined in the Cambrian rocks of the Iberian Peninsula (see Liñán et al., 1993a). They represent the last Lower Cambrian regional stage and the first Middle Cambrian regional stage, respectively. The boundary between these stages coincides with the classical Lower–Middle Cambrian boundary defined in the nineteenth century by the appearance of Paradoxides species, and it closely corresponds with the previous disappearance of archaeocyathids. Although a new Lower–Middle Cambrian boundary has been proposed by Geyer (1990b) and Geyer and Landing (1995, 2004), identified by the appearance of Protolenus (Hupeolenus) species (classic upper Lower Cambrian in the Acadobaltic province), we prefer to maintain the first appearance datum (FAD) of Acadoparadoxides mureroensis as the Lower–Middle Cambrian boundary marker for the Mediterranean subprovince (according to Sdzuy, 1961, 1971a,b, 1995; Liñán and Gozalo, 1986; Liñán et al., 1993a,b, 1996, 2002; Dean and Özgül, 1994; Loi et al., 1995; Pillola et al., 1995; Sdzuy et al., 1996, 1999; Elicki and Pillola, 2004; and Dean, 2005), at least until the International
Commission on Stratigraphy establishes an official subdivision. For a recent discussion, see Sdzuy et al. (1999). The Bilbilian and Leonian Stages were defined in mixed carbonate-siliciclastic facies and used subsequently for both carbonate and siliciclastic sequences in Spain (Sdzuy and Liñán, 1993; Liñán et al. 1995, 2002). These stages were subsequently applied to Sardinia (Loi et al., 1995; Pillola et al., 1995; Perejón et al., 2000; Elicki and Pillola, 2004), Turkey (Dean and Monod, 1997; Dean, 2005), Germany (Elicki, 1997), France (Álvaro et al., 1998b, 2001a,b; Álvaro and Vizcaïno, 2000), Jordan (Rushton and Powell, 1998; Elicki et al., 2002), and Portugal (Liñán et al. 2004a). Both stages are easily identifiable in the Mediterranean subprovince sensu Sdzuy (1971a, 1972) and Sdzuy et al. (1999) because they contain biostratigraphical markers from both inner and outer shelf to basinal environments that can be recognized in Morocco, Spain, France, Italy, Germany, Turkey, Jordan, and Israel. In this paper we compare the lithostratigraphy and biostratigraphy of the Mediterranean subprovince during Bilbilian and Leonian times in order to propose a hypothesis on the correlation of both stages with the main regional stages and levels proposed by the ISCS. BILBILIAN AND LEONIAN STRATIGRAPHY OF THE MEDITERRANEAN SUBPROVINCE The Mediterranean subprovince belongs to the western Gondwana area and corresponds with the so-called European shelf (Fig. 2), in accordance with the paleogeographic proposal of Courjault-Radé et al. (1992). The stratigraphy of the most important areas with Bilbilian and Leonian rocks in the Mediterranean subprovince is shown in Figure 3. The rocks of this time span are mainly siliciclastic in Morocco, Portugal, southern Spain, and Germany; mainly carbonate in northern Spain, France, Italy, and Turkey; and terrestrial with mixed carbonate-siliciclastic marine intercalations in the Middle East (Jordan and Israel). Morocco In the High Atlas and Anti-Atlas, the Bilbilian and Leonian facies are mainly represented by siliciclastic materials (Geyer, 1989; Geyer and Landing, 1995). The top of the Issafen Formation, containing shales and interbedded limestone and belonging to the upper part of the Sectigena zone, may be correlated with
The Lower–Middle Cambrian boundary in the Mediterranean subprovince
Badulesia tenera
Eccaparadoxides sdzuyi Acadoparadoxides mureroensis Protolenus jilocanus
Badulesia tenera
• Ptychagnostus gibbus Kymataspis arenosa
Cephalopyge notabilis
AGDZIAN
? Oryctocephalus indicus • Ovatoryctocara granulata
Protolenus (Hupeolenus)
BANIAN (pars)
Protolenus dimarginatus
Realaspis
MARIANIAN (pars)
FAD
361
CAESARAUGUSTAN (pars)
Ornamentasis frequens
TISSAFINIAN
LEONIAN
Eccaparadoxides asturianus
BILBILIAN
WEST GONWANA ISCS correlation levels COMPOSITE
MOROCCO TOUSHAMIAN (pars)
CAES. (pars)
IBERIA
BANIAN (pars)
Sectigena (pars)
Serrodiscus
Figure 1. Latest Marianian, Bilbilian, Leonian, and earliest Caesaraugustan chronostratigraphic scheme for Iberia, Morocco, and the unified chronostratigraphy proposed by Geyer and Landing (2004). The scheme for Iberia is according to Sdzuy (1971a), Liñán et al. (1993b, 2002) and Dies et al. (2004). The scheme for Morocco is according to Geyer (1990a) and Geyer and Landing (1995). The west Gondwana composite is according to Geyer and Landing (2004). Abbreviations: ISCS—International Subcommission on Cambrian Stratigraphy; FAD—first appearance datum.
• Hebediscus attleborensisCalodiscus-Serrodiscus bellimarginatus-Triangulaspis assemblage
FAD
30° (±5°) NORTH CHINA AND KOREA
Asiatic shelf
BURMA
NEW GUINEA
0° Equator
SOUTH CHINA
E. IRAN
TASMANIA
TIBET
WEST NEW ZEALAND
INDIA
European shelf
EAST EUROPE
ARABIA CENTRAL EUROPE
TURKEY
MADAGASCAR
FRANCE SARDINIA SPAIN
AFRICA
30° (±5°) Americano-African shelf
AUSTRALIA
INDOCHINA
AVALONIA
PIEDMONT
YUCATAN
SOUTH AMERICA
ANTARCTICA
EAST NEW ZEALAND
Figure 2. Paleogeographic reconstruction of the European shelf and location of the areas studied in this article (after Courjault-Radé et al., 1992).
Biozones
Stage
Caesaraugustan (p.)
Par
AO
Issafen Fm (pars)
P
Jbel Wawrmast Fm
Asrir Fm
Tamanart Fm (pars)
Jbel Afraou Fm (pars) B
ANTI-ATLAS
P
Santo Domingo Fm (pars)
Castellar Fm
Par
Es
Archaeocyathans
GERMANY
Par
Tröbitz Fm
B Delitzsch Fm (pars)
Tiefenbach Fm
Galgenberg Fm Par
Triebenreuth Fm
Wildenstein Fm
Doberlug Franc. F.
Pal: Palaeolenus spp. and related genera O: Onaraspis spp. P: First record of Protolenus spp. Cl: Clavigellus spp.
Issafen Fm (pars)
Asrir Fm
A, Pal
Cl
Es
Los Villares Fm (pars)
B
OSSA-MORENA
Daroca Fm
P
Time gap
Nodular carbonate beds (mainly)
San Giovani Fm (pars)
A, Cl P
Campo Pisano Fm (pars)
Carbonate beds (mainly)
B
SARDINIA
Coarse siliciclastic beds (mainly)
Las Tours Fm (upper Mb pars)
Pont du Possarou Fm
La Tanque Fm
Coulouma Fm (pars)
MONTAGNE NOIRE
Mudstone-siltstone-fine sandstone beds (mainly)
Pal
Láncara Fm (lower Mb pars)
O
O
Es
Ea
A, Cl
Es
Ea
B
Oville Fm (pars)
CANTABRIAN
Valdemiedes Fm A
Mansilla Fm
B
Huèrmeda Fm (pars)
B
Murero Fm (pars)
CADENAS IBÉRICAS
Hudai Fm (pars)
A, Cl P
Çal Tepe Fm
Seydisehir Fm (pars)
TAURUS
Zabruck Fm (pars)
Koruk Fm (pars)
B
Sosink Fm (pars)
AMANOS
Pal O
Salib Fm (pars)
Burj Fm
Umm Ishrin Fm (pars)
JORDAN
Figure 3. Generalized correlation table of the Bilbilian, Leonian, and earliest Caesaraugustan stages (late Late Cambrian to early Middle Cambrian) in the Mediterranean subprovince sensu Sdzuy et al. (1999). Abbreviations: FAD—first appearance datum; Fm—Formation; Mb—Member; spp.—species.
Ea: Eccaparadoxides asturianus
Es: Eccaparadoxides sdzuyi
B
Jbel Wawrmast Fm (pars) Ea
HIGH-ATLAS
B: FAD of Badulesia tenera A: FAD of Acadoparadoxides mureroensis Par: First record of Paradoxididae
Realas. FAD
Protolenus Protolenus Acadoparad. Eccaparad. Eccaparad. dimarginatus jilocanus mureroensis sdzuyi asturianus
Badulesia tenera
LEONIAN
m aF car b) Lánpper M (u
BILBILIAN
362 Gozalo et al.
The Lower–Middle Cambrian boundary in the Mediterranean subprovince the lowermost Bilbilian (see Fig. 3 in Geyer and Landing, 2004). The Asrir Formation is usually conformable above the Issafen Formation. It is composed of shales and quartzarenites 30–180 m thick that were deposited in shallow marine environments. Fossils are rare. The lower part of this formation belongs to the Hupeolenus zone (Bilbilian), and the first appearance datum of Acadoparadoxides mureroensis (Leonian) has been reported at the top of it and represents the lower part of the Leonian Stage in both areas (Sdzuy, 1995). The Tamanart Formation occurs above the Asrir Formation in the Amouslek section and laterally changes, in other localities, into the Brèche à Micmacca Member of the Jbel Wawrmast Formation (Geyer and Landing, 1995; Geyer et al., 1995). Abundant trilobites have been recorded in these formations, ranging from the Ornamentaspis frequens zone (Leonian) to the Pardailhania zone (Caesaraugustan). The FAD of Badulesia tenera occurs in the Jbel Wawrmast Formation (High Atlas) and within its lateral equivalent, the Jbel Afraou Formation (Anti-Atlas). This FAD marks the beginning of the Caesaraugustan Stage. The abundant trilobite assemblages found in both regions (Geyer, 1990a, 1994, 1998; Geyer et al., 1995; Sdzuy, 1995; Geyer and Landing, 2004) permit us to recognize the zones of the Bilbilian and Leonian Stages in Morocco. Germany Bilbilian–Leonian rocks have been recognized in two areas: Leipzig and the Franconian Forest (Ludwig, 1969; Sdzuy, 1970; Brause et al., 1997; Elicki, 1997; Geyer and Wiefel, 1997). The Franconian Cambrian of this time interval is represented by a siliciclastic succession and is composed of four formations: Tiefenbach, Galgenberg, Triebenreuth, and Wildenstein (Elicki, 1997). The Tiefenbach Formation consists of over 300 meters (m) of quartzites and siltstones and is supposedly Lower Cambrian or/and lowermost Middle Cambrian in age. The Galgenberg Formation is 40–60 m thick, with shales, medium to fine sandstones, and carbonates; the presence of Middle Cambrian trilobites suggests that it belongs in the Paradoxides insularis zone. The Triebenreuth Formation is 50 m of volcanoclastics rocks with trilobites of the P. oelandicus zone. The Wildenstein Formation is less than 100 m thick and consists of graywackes, sandstones, and shales with carbonate nodules containing trilobites of the P. pinus zone. The last three formations are equivalent in age to the Leonian. A subsurface sequence is known in the Doberlug boreholes (Leipzig area). It is represented by alternating sandstones, quartzites, and shales with rare and very thin carbonate layers. The thickness is more than 500 m (Sdzuy, 1970; Elicki, 1997). The Doberlug IV (pars of the Delitzsch Formation) contains trilobites from the Badulesia tenera zone (Sdzuy et al., 1999). The faunas of the LS 1/63 (pars of the Tröbitz Formation) belongs to the Leonian Stage and includes common species known from Spain, such as Condylopyge cruzensis and Paradoxides? aff. enormis (see Sdzuy, 1970; Dies and Gozalo, 2004).
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Spain Ossa-Morena Zone (South Spain) The Bilbilian and Leonian rocks of the Ossa-Morena zone (southern Spain) crop out as a continuous sequence in the Alconera and Córdoba regions. Relatively good biostratigraphic control is present only in the Córdoba region, where there are several levels with Leonian and Caesaraugustan trilobites in the Los Villares Formation (Liñán Guijarro, 1978; Liñán et al. 1995, 2002; Liñán et al., 2004b). The Lower Bilbilian in the Córdoba region is probably represented by the top of the Santo Domingo Formation in accordance with Liñán et al. (1993a), and it is composed of red shale, dolostone, and chert-bearing limestone with stromatolites, algae, and bioclastic brachiopods. The facies and sedimentological characteristics suggest deposition in supralittoral to restricted infralittoral environments (García Hernández and Liñán, 1983). The Castellar Formation (sensu Liñán et al., 1995) is 75–84 m thick, lies conformably above the Santo Domingo Formation and is composed of sandstone and conglomerate mainly deposited under littoral to shallow sublittoral conditions; it has not yielded fossils of biostratigraphic interest yet. The Los Villares Formation (sensu Liñán et al., 1995) consists of more than 450 m of sandstones and interbedded siltstones. The presence of the trilobites Alueva hastata at the base suggests an age extending from the Leonian (Acadoparadoxides mureroensis zone) through the Caesaraugustan (Badulesia granieri zone) (Liñán Guijarro, 1978; Liñán et al., 1995, 2002, 2004a). Cadenas Ibéricas (Northeast Spain) The Bilbilian–Leonian rocks are mainly represented by the detrital Daroca Formation and mixed facies of the Mesones Group. The first is a terrigenous sequence 90–250 m thick including heterogeneous lithologies (Sdzuy and Liñán, 1993; Álvaro and Vennin, 1998). In some northern localities, trilobites and acritarchs have been found that confirm a Bilbilian age for the Daroca Formation (Gámez et al., 1991; Álvaro and Liñán, 1997; Palacios and Moczydlowska, 1998; Liñán and Gozalo, 2001). The Mesones Group is subdivided into the Valdemiedes, Mansilla, and Murero Formations. It is essentially composed of shales with interbedded carbonate nodules, dolostones, and limestones. It was deposited mainly in sublittoral environments (Sdzuy and Liñán, 1993). The Valdemiedes Formation is 20–250 m thick and is mostly composed of green shales and marly shales with carbonate nodules, as well as scarce and fine carbonate sandstone levels; the position of this formation is lower Bilbilian to middle Leonian (Liñán et al., 2002) based on its diverse record of trilobites (see Liñán and Gozalo, 1986; Gozalo et al., 1993a; Dies Álvarez, 2004); in the upper part of this formation is recorded the Valdemiedes event (Liñán et al., 1993b). The Mansilla Formation, 10–90 m thick, is made up of alternating brown dolostones and limestones and purple and violet shales; trilobites occur in the upper part, including late Leonian to earliest Caesaragustan species (e.g., Gozalo and Liñán, 1995; Sdzuy et al., 1999; Chirivella Martorell et al., 2003). Only the basal levels of the Murero
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Formation correspond to the interval studied herein; they are composed of green lutites with carbonate nodules and interbeds of fine sandstone. The lower boundary is slightly diachronous, yet it coincides approximately with the base of the Caesaraugustan. Cantabrian Mountains (Northwest Spain) In the Cantabrian Mountains, the Láncara and Oville Formations represent the time interval studied here. Fossils of such age have been found only in the Esla nappe area and the Porma section (Zamarreño, 1972; Sdzuy and Liñán, 1993; Sdzuy, 1995; Perejón and Moreno-Eiris, 2003). The Láncara Formation is made up of 150–225 m of dolostone, limestone, and interbedded shale. The lower member includes a persistent dolostone level followed, only in the Esla nappe, by a level of gray lenticular limestones (ooidal grainstones) with archaeocyathans and trilobites (Álvaro et al., 2000b). Biostratigraphic events are recorded only at the top of the limestone level. Lower Cambrian trilobites of the Sectigena zone (Sdzuy, 1995) have been recorded below the Bilbilian archaeocyathans referred from the Valdoré locality (Debrenne and Zamarreño, 1970; Perejón, 1994; Perejón and Moreno-Eiris, 2003), and we have found Palaeolenus sp. in the same level (Fig. 4F–H). The paleontological and sedimentological data suggest supralittoral to littoral environments in the western region and shallow sublittoral conditions in the eastern region (Zamarreño, 1972; Aramburu et al., 1992). The lower part of the upper member of the Láncara
Formation is composed of pink encrinite packstones; the base of this unit is erosive and, in some places, includes pebbles (Álvaro et al., 2000b). In the Valdoré section, these pebbles contain the Lower Cambrian trilobites Kingaspis campbelli (Fig. 4A; Liñán et al., 2003) and Onaraspis sp. (Fig. 4E and I). The earliest specimens of Acadoparadoxides mureroensis have been recorded from the matrix of this level. Just above this pebble-bearing interval, we have found the trilobite Clavigellus sp. (Fig. 4B–D). The upper part of the upper member of the Láncara Formation consists of widespread limestone deposits followed by nodular red limestones (griotte facies). Diachronous boundaries, from Leonian to Caesaraugustan in age (Middle Cambrian), have been inferred for this member based on its trilobite contents (Sdzuy 1968, 1969; Gozalo et al., 1993b; Sdzuy and Liñán, 1993). Recently Geyer and Landing (2004, p. 193–194) discussed the assignation of the Spanish material to Kingaspis campbelli. They noted some small differences, and wrote (p. 194): “One principal character of K. campbelli is that anterior facial suture meets the margin exsagittally posterior to the anterolateral corners of the glabella—a character that is not shown in the Spanish material.” Curiously, in the topotypic material figured by Rushton and Powell (1998, Figs. 21–26) it is possible to observe that the main character noted by Geyer and Landing (2004) is variable between posterior (Fig. 21 in Rushton and Powell, 1998) and anterior (Fig. 24 in Rushton and Powell, 1998) to the anterolateral
C
B D A
F
E H
I
G
Figure 4. Selected trilobites from the Valdoré section (Cantabrian Mountains); bar = 1 mm. (A) Kingaspis (Kingaspis) campbelli (King, 1923); MPZ 2000/8; cranidium preserved in limestones. (B, C, D) Clavigellus sp.; MPZ 2005/296; cranidium preserved in limestones; lateral, frontal, and anterior views. (E) Onaraspis sp.; MPZ 2005/297; incomplete cranidium preserved in limestones. (F) Palaeolenus sp.; MPZ 2005/298; cranidium preserved in limestones. (G) Palaeolenus sp.; MPZ 2005/299; incomplete cranidium preserved in limestones. (H) Palaeolenus sp.; enlarged view of MPZ 2005/299. (I) Onaraspis sp.; enlarged view of MPZ 2005/297.
The Lower–Middle Cambrian boundary in the Mediterranean subprovince corners; thus we considered this character an inner character of the intraspecific variation of Kingaspis campbelli in accordance with the identification of Liñán et al. (2003). The basal levels of the Oville Formation belong to the Leonian or lower Caesaraugustan in some areas; they are mainly glauconitic sandstones and green shales, with some carbonate nodules. The lowest part of the formation contains several fossiliferous levels with trilobites (Sdzuy, 1968, 1969). Diachroneity between this formation and the preceding formation is shown by the presence of different trilobite assemblages ranging from upper Leonian to upper Caesaraugustan in age (Zamarreño, 1972; Sdzuy and Liñán, 1993; Liñán et al., 2002). France The successions of uppermost Marianian or Bilbilian to basal Caesaraugustan rocks in the Montagne Noire comprise four formations: the upper member of Las Tours, Pont du Poussarou, La Tanque, and lowest Coulouma (Álvaro et al., 1998b, 2001a,b). The upper member of the Las Tours Formation is 40–150 m thick; thin-bedded limestones are intercalated within green shales. Ferralsia blayaci has been recorded in this member and indicates either a late Marianian or an earliest Bilbilian age (Álvaro et al., 1998a). The Pont du Poussarou Formation is a 20–80 m massive limestone with dolomitic intercalations; it is probably Bilbilian in age. The La Tanque Formation reaches 60 m of reddish and purple carbonates and interbedded shales; the age is probably Leonian. The Pont du Poussarou and the La Tanque Formations lack biostratigraphic markers; thus their ages are inferred from their stratigraphic positions in the succession. The base of the Coulouma Formation (green shales with carbonate nodules) contains the trilobite Asturiaspis inopinatus, of upper Leonian age, and Badulesia tenera, which marks the base of the Caesaraugustan Stage (see Álvaro and Vizcaïno, 2000). Italy The Bilbilian and Leonian of Sardinia (Italy) are represented by two carbonate formations, San Giovanni and Campo Pisano (Pillola, 1991; Loi et al. 1995; Pillola et al., 1995; Perejón et al., 2000; Elicki and Pillola, 2004). The first is composed mainly of massive black to gray limestones over 300 m thick, representing mostly a shallow-water carbonate succession. Only rare archaeocyathans have been recorded from the upper part of the San Giovanni Formation, which has been assigned to the Toyonian 2–3 (Debrenne and Gandin, 1985; Zhuravlev, 1995). It is overlain by slightly deeper sediments of the Campo Pisano Formation, mostly 40–60 m thick (well-bedded and nodular limestones, with calcareous shales and marls). This unit represents a condensed succession, spanning the uppermost Lower Cambrian (Bilbilian) through the Middle Cambrian (Caesaraugustan). At the base of the Campo Pisano Formation, the oldest assemblage, CP1, contains Protolenus pisidianus and Clavigellus? n. sp. (Loi et al., 1995; Elicki and Pillola, 2004); this assemblage is considered
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uppermost Lower Cambrian. A few meters above, in the lower part of assemblage CP2, lies the FAD of Acadoparadoxides mureroensis (Leonian). The upper part of level CP2 contains Pardailhania hispida, which indicates the Caesaraugustan Stage. Turkey In Turkey, Bilbilian, Leonian, and Caesaraugustan levels have been dated. The formations on which Lower and Middle Cambrian strata rest vary considerably. For example, the oldest “red nodular facies” are dated as basal Leonian and occur in the western Taurides (Çal Tepe Formation); the youngest and more southernly occurrences in the Border Folds (Koruk Formation) belong to the middle or upper Caesaraugustan (Dean and Monod, 1997). The two sequences that have been dated by means of Bilbilian, Leonian, and basal Caesaraugustan trilobites are found in the Hüdai and Çal Tepe areas (western Taurides: Dean and Özgül, 1994; Dean, 2005) and in Alan Yayla (Amanos Mountains: Dean et al., 1986). Cambrian sequences from the Taurus Mountains comprise the Hüdai Quartzite, Çal Tepe, and Seydisehir Formations (Dean and Monod, 1970; Dean and Özgül, 1981, 1994; Dean, 2005). The Hüdai Quartzite Formation consists of less than 500 m of thick-bedded, reddish white and gray, laminated quartz-arenite, which exhibits cross-bedding and is considered to represent a beach deposit. The Çal Tepe Formation, which is 130–170 m thick, has been subdivided into four carbonate members in its type area (e.g., Dean, 2005): the dolomite member, 80–150 m, is composed mainly of thickly bedded, coarsely crystalline, brown dolostone; the black limestone member, over 20 m thick, comprises thickly bedded, tough, dark gray to black limestone indicative of a shallow, open marine environment; the light gray limestone member comprises little more than 10 m of pale, thinly bedded limestones, indicating a high-energy, shallow marine environment; finally, the red nodular limestone member comprises 40–16 m of medium-bedded, pink beige and light gray nodular limestone, often with intercalations of bright red mudstone, and marks a transition to lower-energy, deeper marine conditions. The overlying Seydisehir Formation is mainly composed by green to gray shales with thin levels of fine-grained, quartzitic sandstone interbedded. Fossils have only been found in the three last members of the Çal Tepe Formation in the Formation type area and near Hüdai. The trilobite species Protolenus pisidianus, Acadoparadoxides mureroensis, and Clavigellus venustus have been found in the light gray limestone member and permitted recognition of the base of the Leonian Stage in this area (Dean and Özgül, 1994; Dean, 2005). In other localities of the Taurus Mountains, the higher red nodular limestones member of the Çal Tepe Formation and the lowest meters of the succeeding siliciclastics of the Seydisehir Formation also contain Pardailhania sp. (Dean and Özgül, 1981; Dean and Monod, 1997; Dean, 2005), which marks the Caesaraugustan Stage. In the Amanos Mountains, the Cambrian sequence comprises the Zabuk, Koruk, and Sosink Formations (Dean et al.,
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Gozalo et al.
1986; Dean and Monod, 1997). The Zabuck Formation is ~100 m thick and consists of thickly bedded, pink, fluviatile or deltaic arkoses. The Koruk Formation consists mostly of 185 m of dark, thickly bedded dolostones or dolomitic limestones with thin shale interbeds; the succession begins with coarse to fine-grained, gray dolostones that are interpreted as typical supratidal to intertidal deposits initiating a transgressive sequence. In some sections the dolostone member changes conformably upward into black, bioturbated packstones that form the lowest part of the limestone member and are interpreted as a progressive change to shallow, open marine environments; the upper part of the formation is composed of over 10 m of gray and red nodular limestones. The Sosink Formation is 250 m of calcareous shales and shally limestones. Badulesia tenera has been found in the lower part of the Sosink Formation near the Alan Yayla (Dean and Krummenacher, 1961; Dean et al., 1986), indicating the beginning of the Caesaraugustan Stage. In other localities, the red nodular limestones on the top of the Koruk Formation have been found to contain middle Caesaraugustan trilobites (Dean and Monod, 1997), showing an evident diachrony in the boundary between the Koruk and the Sosink Formations. Jordan and Israel The Cambrian succession in southern Israel was divided by Soudry and Weissbrod (1995) into four formations (Amudei Shelomo, Timna, Shehoret, and Netafim). The lowest and uppermost formations are fluvial deposits, while the Timna Formation accumulated in marine depositional environments; it marks the earliest Paleozoic transgression of the paleo-Tethys on the Arabian-African craton. It contains some carbonate levels in a mainly siliciclastic succession and includes datable fossils, ichnofossils, and stromatolitic structures representing shallow-water and intertidal environments in a marginal basin (Soudry and Weissbrod, 1995, p. 340). The sequence from the Dead Sea is composed of three formations (Salib, Burj, and Umm Ishrin). As in Israel, the lower and upper ones are fluvial units, while the mixed carbonate-siliciclastic Burj Formation shows, from its base upward, a clear pattern of marine transgression and regression (Rushton and Powell, 1998). This formation has been divided into three members, from the bottom the Tayan Member (up to 20 m of greenish to reddish siltstone and fine-grained sandstone), the Numayri Member (50–60 m of limestone and dolostone), and the Hanneh Member (~30 m of greenish to reddish sandstone and minor siltstone) sensu Elicki et al. (2002). The maximum transgressive phase produced shallow subtidal to supratidal environments at the intermediate Numayri Member. Trilobite faunas are recorded from the Timna and Burj Formations (see Parnes, 1971; Rushton and Powell, 1998). These formations are considered equivalent by Rushton and Powell (1998). Trilobite assemblages of the Burj Formation studied by the latter authors included a level with Kingaspis campbelli and Palaeolenus antiquus at the top of the Numayri Member and,
slightly below, Realaspis sp. nov., and Redlichops blanckenhorni. Rushton and Powell also found Onaraspis palmeri in this member; this species was defined in the Timna Formation from levels correlated with Kingaspis campbelli (see Parnes, 1971). In summary, the Bilbilian and Leonian facies distribution in the Mediterranean subprovince (Fig. 3) is siliciclastic in Morocco, southern Spain, and Doberlug and may represent siliciclastic segments of an old complex marine shelf. Northern Spain, France, Sardinia, and Turkey may represent the carbonate segments of the same marine shelf during Leonian time. The Middle East represents a cratonic sequence. Mixed facies from the Cadenas Ibéricas were deposited in a marine gulf (Sdzuy and Liñán, 1993). BILBILIAN BIOCHRONOLOGY The Bilbilian is considered to represent the uppermost Lower Cambrian in Spain; it contains easily recognizable trilobites and acritarchs (Liñán et al., 2002), and it can be correlated with accuracy to other areas of the Mediterranean subprovince. Its upper limit is defined by the top of the Protolenus jilocanus interval zone. However, its lower boundary is not well characterized. Liñán et al. (1993a, p. 827) proposed the regional last record of the genera Andalusiana and Serrodiscus as the upper boundary of the previous Marianian stage; it has also been considered the FAD of the genus Realaspis (Liñán et al., 2002). Sdzuy (1971a) coined the term “Bilbilian” to define a mixed siliciclastic-carbonate sequence without Olenellidae or Paradoxididae but with the presence of Protolenidae, Ellipsocephalidae, Redlichiidae, and genera similar to Onaraspis. Sdzuy (1971a) and Liñán et al. (1993a) choose the Huérmeda section, in the Jalón Valley (Zaragoza province) near Huérmeda and Calatayud (which are near the old Roman city of Bilbilis), as the stratotype section for the Bilbilian Stage. In the Cadenas Ibéricas this stage includes strata from shales of the Huérmeda Formation to the last Lower Cambrian rocks within the Valdemiedes Formation. In the Cantabrian region, the Bilbilian is represented by archaeocyath-bearing limestones of the Valdoré Formation (Debrenne and Zamarreño, 1970). Liñán et al. (1993a), in the original definition of the Bilbilian Stage, proposed the Valdoré section as a reference section in the Cantabrian Mountains. Although the taxa described in the first works about the Bilbilian Stage were mainly endemic forms, subsequent works have recognized several species defined in other regions of the Mediterranean subprovince and vice versa. Furthermore, some genera originally described in distant biogeographical regions have also been described in the Bilbilian of Spain. These finds have allowed the establishment of reliable correlations between the Lower–Middle Cambrian stratigraphic units of different Mediterranean areas: Jordan and Israel (Parnes, 1971; Rushton and Powell, 1998), Turkey (Dean and Özgül, 1994; Dean and Monod, 1997), Morocco (Sdzuy, 1995; Sdzuy et al., 1999), Italy (Pillola et al., 1995; Perejón et al., 2000), Germany (Sdzuy, 1971a; Sdzuy et al., 1999), and Spain (Liñán et al., 2002). Furthermore, the presence of the genus Onaraspis in the Cadenas Ibéricas, the
The Lower–Middle Cambrian boundary in the Mediterranean subprovince Cantabrian Mountains, Morocco, Israel, and Jordan allows good correlation to Australia (see Gozalo and Liñán, 1997; Geyer and Landing, 2004). Finally, the presence of Palaeolenus sp. together with archaeocyathans in the Cantabrian Mountains can be a useful tool for correlation with the last Lower Cambrian archaeocyathan levels in Sardinia, China, Australia, and Siberia (see Zhuravlev, 1995). The recent revision of the taxonomy of Bilbilian trilobites (Dies Álvarez, 2004) allows proposal of the following zonation for the upper part of this stage in the Cadenas Ibéricas. The main problem is that uppermost Marianian and lowermost Bilbilian strata in the Cadenas Ibéricas lack fossils. Thus the lowest levels of the Bilbilian Stage remain without precise faunal zonation. According to Sdzuy (1971a) and Liñán et al. (1993a), the first levels of the Bilbilian Stage, which contain Realaspis but not Protolenus (Hupeolenus), are found near Los Cortijos de Malagón (central Spain). This fauna could be correlated with the top of the Sectigena zone in Morocco (Banian Stage). Realaspis has also been recorded in Jordan (Rushton and Powell, 1998) and may be the ancestor of the genus Onaraspis. Protolenus dimarginatus zone: This is an interval biozone with the lower boundary determined by the FAD of P. dimarginatus. The upper boundary is placed at the FAD of P. jilocanus. The FAD of P. jilocanus coincides with the FAD of Hamatolenus (Hamatolenus) ibericus (see Liñán et al., 1993b; Dies et al., 2004), so this species can determine the upper boundary of the zone if P. jilocanus is not present. In the Cadenas Ibéricas, this zone contains P. dimarginatus (Fig. 5A), Kingaspis (K.) campbelli (Fig. 5F), and P. termierelloides (Fig. 5G). Protolenus jilocanus zone ( = Hamatolenus (H.) ibericus sensu Liñán et al., 1993b). The original definition of the zone is confirmed with very few changes as follows (Dies Álvarez, 2004): An interval zone with its lower boundary placed at the FAD of P. jilocanus and/or of H. (H.) ibericus; the upper boundary is marked by the FAD of Acadoparadoxides mureroensis. The trilobite species recorded in the P. jilocanus zone of the Cadenas Ibéricas include H. (H.) ibericus (Fig. 5D), H. (Myopsolenus) sp. A, Protolenus jilocanus (Fig. 5E), P. termierelloides, P. pisidianus (Fig. 5B), P. interscriptus, Kingaspis (K.) campbelli, Sdzuyia sanmamesi, Tonkinella sequei, Onaraspis altus (Fig. 5C), and Alueva undulata. Protolenus dimarginatus, P. interscriptus, and P. termierelloides were defined in Morocco by Geyer (1990a) in the Hupeolenus zone. Thus the presence of these species in both zones allows a precise correlation between Morocco and the Cadenas Ibéricas. The stratotype of the Bilbilian-Leonian boundary is located in Murero (Cadenas Ibéricas), within the upper Valdemiedes Formation. Liñán et al. (1993b) identified the Valdemiedes event in the uppermost Bilbilian rocks. The event was identified by means of changes in sediment mineralogy and trace fossil contents, as well as the disappearance of numerous trilobite species and dwarfing of the brachiopod fauna. The Valdemiedes event is recorded, at its type locality, Murero, within a conformable, expanded, monofacial succession (in a transgressive context),
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consisting of mudstones and siltstones that (according to illite crystallinity values) did not surpass the anchizone. Furthermore, some authors used outcrops of the upper Valdemiedes Formation for studies on cyclostratigraphy without finding any important sedimentary break (Álvaro et al., 2000a). The Valdemiedes event has been identified in all Lower or Middle Cambrian localities of Cadenas Ibéricas (northeast Spain). The paleogeographic extent of the Valdemiedes event (and thus its probable significance as a geoevent) is a matter for further research, but its chronological position suggests that it was close to or coincided with the widely accepted Lower or Middle Cambrian mass extinction. The Valdemiedes event may be recognized only in areas where uppermost Lower Cambrian sublittoral marine ecosystems were reestablished after the early Bilbilian Daroca regression or its equivalents (such as the Hawke Bay regression). In areas where the Lower–Middle Cambrian transition produced coarse clastic, the Valdemiedes event may be difficult to discriminate from the Daroca regression. In areas where late Early Cambrian regression produced a hiatus extending into the Middle Cambrian (as, for example, in Scandinavia), the Valdemiedes event may be included in this gap (Liñán et al., 2002). LEONIAN BIOCHRONOLOGY The lower boundary of the Leonian Stage was placed at the FAD of Acadoparadoxides mureroensis (see Fig. 5H–J). This FAD is one of the most widespread in the Middle Cambrian, for it is recorded in the Cantabrian Mountains and the Cadenas Ibéricas (Spain), Sardinia (Italy), the High Atlas (Morocco), the Taurus Mountains (Turkey), Tuva (Russia), and perhaps also Newfoundland (Canada); see Sdzuy et al. (1999). The Leonian upper boundary has been placed at the FAD of Badulesia tenera, which has been recorded from Rhode Island (USA), New Brunswick and Newfoundland (Canada), the High Atlas (Morocco), Sierra Morena, the Cantabrian Mountains, and the Cadenas Ibéricas (Spain), Doberlug (Germany), the Amanos Mountains (Turkey), and the Montagne Noire (France). The sequences from the Cadenas Ibéricas (mixed facies) and the Cantabrian Mountains (carbonate facies) are some of the most continuous and fossiliferous successions known from the Leonian Stage. They permit identification of three trilobite zones named the Acadoparadoxides mureroensis zone, the Eccaparadoxides sdzuyi zone, and the Eccaparadoxides asturianus zone (Gozalo and Liñán, 1995; Sdzuy et al., 1996, 1999). These are interval zones limited by the first appearance of the Paradoxides species chosen from a phylogenetic line, and they are also characterized by other trilobite taxa. The lower boundary of the Acadoparadoxides mureroensis zone is marked by the FAD of the trilobite species A. mureroensis; the position has been considered the Lower–Middle Cambrian boundary in different regions (Sdzuy, 1971b, 1972, 1995; Álvaro et al., 1993; Gozalo et al., 1993b; Liñán et al., 1993a,b; Dean and Özgül, 1994; Loi et al., 1995; Pillola et al., 1995; Elicki and
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Gozalo et al.
B
A D
C
E
F
G
H
I
J
Figure 5. Selected trilobites from sections in Cadenas Ibéricas. (A) Protolenus dimarginatus Geyer, 1990a; MPZ 01/100; internal mold of cranidium preserved in shales; Jarque 1 section; bar = 2 mm. (B) Protolenus pisidianus Dean in Dean and Özgül, 1994; MPZ 2004/89; cranidium preserved in limestones; Ateca 16 section; bar = 2 mm. (C) Onaraspis altus (Liñán and Gozalo, 1986); MPZ 99/515 and MPZ 99/586; internal molds of complete specimens preserved in shales; Rambla de Valdemiedes 1 section; bar = 2 mm. (D) Hamatolenus (Hamatolenus) ibericus Sdzuy, 1958; MPZ 99/184; internal mold of a complete specimen preserved in shales; Rambla de Valdemiedes 1 section; bar = 5 mm. (E) Protolenus jilocanus (Liñán and Gozalo, 1986); MPZ 01/107; internal mold of cranidium preserved in shales; Jarque 1 section; bar = 2mm. (F) Kingaspis campbelli (King, 1923); MPZ 99/34; internal mold of cranidium preserved in shales; Jarque 1 section; bar = 2 mm. (G) Protolenus termierelloides Geyer, 1990a; MPZ 2004/85; incomplete cranidium preserved in limestones; Ateca 16 section; bar = 2 mm. (H) Acadoparadoxides mureroensis (Sdzuy, 1958); MPZ 01/84; internal mold of pygidium (morphotype A) preserved in shales; Rambla de Valdemiedes 1 section; bar = 2 mm. (I) Acadoparadoxides mureroensis (Sdzuy, 1958); MPZ 01/ 83; internal mold of pygidium (morphotype B) preserved in shales; Rambla de Valdemiedes 1 section; bar = 2 mm. (J) Acadoparadoxides mureroensis (Sdzuy, 1958); MPZ 01/82; internal mold of incomplete specimen preserved in shales; Rambla de Valdemiedes 1 section; bar = 2 mm.
The Lower–Middle Cambrian boundary in the Mediterranean subprovince Pillola, 2004). Its upper boundary is the FAD of Eccaparadoxides sdzuyi. Acadoparadoxides mureroensis is the guide fossil of this zone and disappears at the base of the following zone. The trilobite species Hydrocephalus cf. harlani, Alueva hastata, Alueva moratrix, and Macannaia are also typical of this interval zone. Other bioevents are the FAD of Condylopyge cruzensis and the last records of the genera Hamatolenus, Protolenus, Latoucheia, and Alueva. The following Eccaparadoxides sdzuyi zone has its upper boundary immediately below the FAD of Eccaparadoxides asturianus. The species Conocoryphe (Conocoryphe) ovata, Acadolenus decorus, and Peronopsella prokovskajae prokovskajae are typical of this zone. Other important bioevents are the FADs of the genera Conocoryphe, Cornucoryphe, Bailiella, Asturiaspis, Skreiaspis, Acadolenus, Peronopsella, and Dawsonia. The upper boundary of the Eccaparadoxides asturianus zone is marked by the FAD of Badulesia tenera. At present, the trilobite species Acadolenus inornatus, Jincella? sulcata, Conocoryphe (Parabailiella) sebarensis, Skreiaspis tosali, Holocephalina? leve, and Bailiaspis dalmani are exclusive to this zone. The FADs of Peronopsella prokovskajae ovetense, Eccaparadoxides sulcatus, Conocoryphe (P.) languedocensis, and Bailiaspis cf. tuberculata are also included in the zone. GLOBAL CORRELATION In recent years, the time interval revisited in this work has been discussed in several papers, and several correlation charts have been produced (i.e., Sdzuy et al., 1999; Geyer and Shergold, 2000; Fletcher, 2003; Geyer and Landing, 2004; Fletcher et al., 2005). There is no general agreement on the main levels. For this reason, we have tried herein to obtain a general view for the Mediterranean subprovince. This allows for a better understanding of regional correlation and for improved global correlation. The trilobite markers of the Bilbilian and Leonian biozones and their respective associations permit fairly precise global correlation of the Bilbilian and Leonian Stages. The main problem is with the lower Bilbilian, because, for the moment, it lacks a biozonation and its lower boundary is not well established. The base of the Bilbilian Stage is defined by the FAD of the genus Realaspis, which has been found in central Spain (Sdzuy, 1961, 1971a) and Jordan (Rushton and Powell, 1998). The next biostratigraphic level is the FAD of Protolenus. Perhaps this level is the best way to recognize this stage. Protolenus shows a wide distribution in the Acadobaltic province, and allows recognition of the last stage of the Lower Cambrian without Paradoxididae. Other important trilobite taxa for correlation are Kingaspis campbelli (see Rushton and Powell, 1998; Liñán et al., 2003), Onaraspis (Gozalo and Liñán, 1997; Geyer and Landing, 2004), and Palaeolenus. Geyer and Landing (2004) proposed the Onaraspis clade as a useful correlation tool, including the genera Onaraspis and Myopsolenites. We believe that the second is a junior synonym of Onaraspis; actually, the name Myopsolenites was a misprint of Öpik (1975), as Rushton and Powell (1998) discussed.
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On the other hand, the tenth correlative unit for the Early Cambrian proposed by Zhuravlev (1995), which corresponds to the mid-Toyonian Lermontovia grandis trilobite zone, is equivalent to the archaeocyathan Irinaecyathus schabanovi-Archaeocyathus okulitchi zone of the Siberian Platform. This unit is identifiable by its trilobite and archaeocyathan assemblages. The archaeocyathan assemblage has been recorded in the Siberian Platform, the Altay-Sayan foldbelt, Australia, China, Spain, and Sardinia (see Zhuravlev, 1995). As for Spain, this archaeocyathan assemblage has been found only in the Bilbilian of the Cantabrian Mountains (Debrenne and Zamarreño, 1970; Perejón and Moreno-Eiris, 2003). Now we report the finding of these archaeocyathans together with the trilobite Palaeolenus sp. This fact reinforces the correlation to China, where the archaeocyathans belong to the Megapaleolenus zone of the Tsanglangpuan Stage (Yuan et al., 2001). Recently Lin and Peng (2004) reviewed the genus Palaeolenus and considered the genus Megapaleolenus a junior synonym of Palaeolenus; furthermore, they agreed with Rushton and Powell (1998), who transferred Schistocephalus antiquus to Palaeolenus. They also commented that “Gigoutella is more likely also a junior synonym of Palaeolenus.” We are in agreement with their systematic opinions. The boundary between the Bilbilian and the Leonian is placed at the FAD of Acadoparadoxides mureroensis. This FAD is close to the FADs of Palaeolenus antiquus, Macannaia spp., and Ovatoryctocara granulata, which may be considered one of the best species for global correlation. The FADs of Badulesia tenera, Parasolenopleura aculeata, Ctenocephalus (Hartella) spp., and Ptychagnostus gibbus may be considered the characteristic bioevents for global correlation of the Leonian upper boundary. A tentative correlation chart of the Bilbilian and Leonian Stages is presented in Figure 6. Bilbilian correlation is based on the presence of the genera Realaspis, Protolenus, and Onaraspis and some species such as Palaeolenus antiquus and Kingaspis campbelli, as well as on the last archaeocyathans. Palaeolenus antiquus was defined in Siberia, and its FAD is considered a marker of the Amgan Stage; this species is recorded slightly below the first Paradoxides (see Korovnikov, 2001, 2004). The lower boundary of the Leonian Stage is based on the association of Acadoparadoxides mureroensis and Macannaia in Spain. Acadoparadoxides mureroensis is also recorded in Morocco, Siberia, Sardinia, and Turkey, and probably in Newfoundland, while Macannaia is present in Laurentia, Australia, and Siberia. Another important taxon is Hydrocephalus harlani, originally defined in Massachusetts (USA), which also occurs in Morocco and probably in Spain (Geyer and Landing, 2001; Dies Álvarez, 2004; Fletcher et al., 2005). Finally, Tonkinella aff. breviceps is recorded from the upper Leonian and has been correlated with the Hsuchuangian Stage (North China) and the Ehmaniella Zone (USA) by Gozalo et al. (2003). The upper boundary of the Leonian is based on the presence of Badulesia tenera and Ptychagnostus gibbus, but they have not been found to co-occur yet. Badulesia tenera is recorded in
Eccaparadoxides bennetti
Eccaparadoxides sdzuyi
Realaspis
Callavia
Protolenus
Cephalopyge
Kiskinella cristata
Holmia kjerulfi
“Ornamentaspis” linnarssoni
“Protolenus”
Paradoxides insularis
Paradoxides pinus
Ptychagnostus gibbus
BALTICA
Bergeroniaspis ornata
Bergeroniellus ketemensis
Lermontovia grandis
Anabaraspis splendens
Oryctocara/ Schistocephalus antiquus
Kounamkites
Ptychagnostus gibbus
SIBERIA
Pararaia bunyerooensis
Pararaia janeae
Xystridura templetonensis / Redlichia chinensis
Ptychagnostus gibbus
AUSTRALIA
Drepanuroides
Palaeolenus
Megapalaeolenus
Redlichia chinensis
Redlichia nobilis
Yaojiayuella
Shantungaspis
Poriagraulos Hsuchuangia Ruichengella
NORTH CHINA
Arthricocephalites/ Changaspis
Protoryctocephalus
Bathynotus
Oryctocephalus indicus
Ptychagnostus gibbus
SOUTH CHINA
O. i. FAD
Olenellus / Bonnia
Eokochaspis nodosa
Amecephalus O. g. FAD arrojoensis
Oryctocephalus indicus
Albertella
Ehmaniella
Bolaspidella
LAURENTIA
Figure 6. Global correlation of the latest Marianian, Bilbilian, and Leonian and earliest Caesaraugustan stages with other regional stages and zones, based on Sdzuy et al. (1999), Geyer and Shergold (2000), Peng and Babcock (2001), Peng (2003), Geyer and Landing (2004), and Fletcher et al. (2005). The scheme for Laurentia is according to the latest modification by Sundberg and McCollum (2003) and Sundberg (2005). Abbreviations: FAD—first appearance datum; O. i. FAD—FAD of Oryctocephalus indicus; O. g. FAD—FAD of Ovatoryctocara granulata.
Serrodiscus
FAD
FAD
Protolenus dimarginatus
Protolenus jilocanus
Acadoparadoxides mureroensis
Agraulos affinis
Eccaparadoxides asturianus
Hydrocephalus harlani
Badulesia tenera
Badulesia tenera
AVALONIA P. paradoxissimus Paradoxides oelandicus
IBERIA Late TEMPLETONIAN (pars)
Early TEMPLETONIAN/ORDIAN
CAESARAUGUSTAN (p)
LEONIAN
BILBILIAN
MARIANIAN (p)
AMGAN TOYONIAN BOTOMAN (pars)
TAINIANGIAN DUYUNIAN
HSUCHUANGIAN (p) MAOCHUANGIAN LUNGWANGMIAOAN TSANGLANGPUAN (pars)
TOPAZAN (pars) DELAMARAN DYERAN (pars)
370 Gozalo et al.
The Lower–Middle Cambrian boundary in the Mediterranean subprovince
Spain, Morocco, and Newfoundland, while Ptychagnostus gibbus has been identified from Laurentia, North China, Australia, Siberia, and Baltica. This correlation is supported by the association of Ctenocephalus (Hartella) and Parasolenopleura aculeata with Badulesia tenera in Spain and Newfoundland (Liñán et al., 1995; Fletcher, 2005) and with Ptychagnostus gibbus in Baltica. In conclusion, the Bilbilian and Leonian Stages are useful Lower–Middle Cambrian stages for the Mediterranean subprovince and also represent a good tool for correlation with other regional stages defined in the upper Lower and lower Middle Cambrian. The Bilbilian–Leonian boundary roughly coincides with the FAD of Ovatoryctocara granulata, one of the levels studied by the ISCS. The Oryctocephalus indicus level is probably located slightly above the Bilbilian–Leonian boundary, and it correlates to somewhere within the lowermost part of the Acadoparadoxides mureroensis zone (see Fig. 6). The close and solomonic occurrence of the Bilbilian–Leonian boundary between the two Global Standard Stratotype-section and Points proposed by the ISCS and the absence of Oryctocara granulata and Oryctocephalus indicus in the Mediterranean region made this classic Cambrian Series boundary a neccessary and complementary reference for intercontinental correlation. Finally, the Bilbilian–Leonian boundary is placed after the global lowermost Cambrian mass extintion event and before the great Middle Cambrian radiation event in the Earth’s History. ACKNOWLEDGMENTS We thank Dr. Loren Babcock (Ohio State University) and Dr. Olaf Elicki (Freiberg University) for their comments. We acknowledge financial support from the Spanish Dirección General de Investigación and FEDER (Projects BTE2003–04997 and CGL2006-12975/BTE) and from Gobierno de Aragón, Museo Paleontológico Group. M.E. Dies Álvarez is a postdoctoral researcher (Ref. Ex 2005-1019) of the Spanish Ministry of Sciences at Lund University. REFERENCES CITED Álvaro, J.J., and Liñán, E., 1997, Nuevos datos acerca del Bilbiliense (Cámbrico Inferior terminal) en las Cadenas Ibéricas y su correlación con otras áreas: Revista Española de Paleontología, v. 12, p. 277–280. Álvaro, J., and Vennin, E., 1998, Stratigraphic signature of a terminal Early Cambrian regressive event in the Iberian Peninsula: Canadian Journal of Earth Sciences, v. 35, p. 402–411, doi: 10.1139/cjes-35-4-402. Álvaro, J., and Vizcaïno, D., 2000, Nouvel assemblage des trilobites dans le Cambrien moyen de la nappe de Pardailhan (Montagne Noire, France): Implications biostratigraphiques dans la région méditerranéenne: Eclogae Geologicae Helvetiae, v. 93, p. 277–289. Álvaro, J., Gozalo, R., Liñán, E., and Sdzuy, K., 1993, The palaeogeography of the northern Iberia at the Lower–Middle Cambrian transition: Bulletin de la Société Géologique de France, v. 164, no. 6, p. 843–850. Álvaro, J.J., Liñán, E., and Vizcaïno, D., 1998a, Biostratigraphical significance of the genus Ferralsia (Lower Cambrian, Trilobita): Geobios, v. 31, p. 499–504, doi: 10.1016/S0016-6995(98)80121-6. Álvaro, J.J., Courjault-Radé, P., Chauvel, J.J., Dabard, M.P., Debrenne, F., Feist, R., Pillola, G.L., Vennin, E., and Vizcaïno, D., 1998b, Nouveau découpage stratigraphique des séries cambriennes des nappes de Pardailhan et
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du Minervois (versant sud de la Montagne Noire): Géologie de la France, v. 1998, no. 2, p. 3–12. Álvaro, J.J., Vennin, E., Muñoz, A., Sánchez-Valverde, B., and Ojeda, J.L., 2000a, Interplay of orbital forcing and tectonic pulses in the Cambrian Iberian platform, NE Spain: International Journal of Earth Sciences, v. 89, p. 366–376. Álvaro, J.J., Vennin, E., Moreno-Eiris, E., Perejón, A., and Bechstädt, T., 2000b, Sedimentary patterns across the Lower–Middle Cambrian transition in the Esla nappe (Cantabrian Mountains, northern Spain): Sedimentary Geology, v. 137, nos. 1–2, p. 43–61, doi: 10.1016/S0037-0738(00)00134-2. Álvaro, J., Debrenne, F., and Vizcaïno, D., 2001a, The Lower Cambrian of the southern Montagne Noire: Annales Société Géologique du Nord, ser. 2, v. 8, no. 4, p. 201–204. Álvaro, J., Lefebvre, B., Shergold, J.H., and Vizcaïno, D., 2001b, The Middle–Upper Cambrian of the southern Montagne Noire: Annales Société Géologique du Nord, ser. 2, v. 8, no. 4, p. 205–211. Aramburu, C., Truyols, J., Arbizu, M., Méndez-Bedia, I., Zamarreño, I., GarcíaRamos, J.C., Suárez de Centi, C., and Valenzuela, M., 1992, El Paleozoico Inferior de la Zona Cantábrica, in Gutiérrez-Marco, J.C., Saavedra, J., and Rábano, I., eds., Paleozoico Inferior de Ibero-América: Badajoz–Cáceres, Spain, Universidad de Extremadura, p. 397–421. Babcock, L.E., Peng, S.C., Geyer, G., and Shergold, J.H., 2005, Changing perspectives on Cambrian chronostratigraphy and progress toward subdivision of the Cambrian System: Geosciences Journal, v. 9, p. 101–106. Brause, H., Elicki, O., and Kemnitz, H., 1997, Lausitzer Masse: Courier Forschungsinstitut Senckenberg, v. 200, p. 298–334. Chirivella Martorell, J.B., Gozalo, R., and Liñán, E., 2003, Estudio taxonómico y biométrico de Badulesia tenera (Hartt in Dawson, 1868), TRILOBITA: Caesaraugustiense (Cámbrico Medio) de las Cadenas Ibéricas (NE de España): Revista Española de Paleontología, v. 18, p. 83–101. Courjault-Radé, P., Debrenne, F., and Gandin, A., 1992, Paleogeographic and geodynamic evolution of the Gondwana continental margins during the Cambrian: Terra Nova, v. 4, p. 657–667. Dean, W.T., 2005, Trilobites from Çal Tepe Formation (Cambrian), near Seydisehir, Central Taurides, Southwestern Turkey: Turkish Journal of Earth Sciences, v. 14, p. 1–71. Dean, W.T., and Krummenacher, R., 1961, Cambrian trilobites from the Amanos Mountains, Turkey: Palaeontology, v. 4, no. 1, p. 71–81. Dean, W.T., and Monod, O., 1970, The Lower Palaeozoic stratigraphy and faunas of the Taurus Mountains near Beysehir, Turkey, I: Stratigraphy: Bulletin of the British Museum (Natural History): Geology, v. 19, no. 8, p. 413–426. Dean, W.T., and Monod, O., 1997, Cambrian development of the Gondwanaland margin in southeastern Turkey: Turkish Association of Petroleum Geologists Special Publication 3, p. 61–74. Dean, W.T., and Özgül, N., 1981, Middle Cambrian trilobite succession in the Çaltepe Formation at Bagbasi (Hadim-Konya) Central Taurus, Turkey: Bulletin of the Mineral Research and Exploration Institute of Turkey, v. 92, p. 1–6. Dean, W.T., and Özgül, N., 1994, Cambrian rocks and faunas, Hüdai area, Taurus Mountains, southwestern Turkey: Bulletin de l’Institut Royal des Sciences Naturelles de Belgique, Sciences de la Terre, v. 64, p. 5–20. Dean, W.T., Monod, O., and Günay, Y., 1986, Lower Palaeozoic stratigraphy in the Southern and Central Amanos Mountains, south central Turkey: Geological Magazine, v. 123, no. 3, p. 215–226. Debrenne, F., and Gandin, A., 1985, La formation de Gonnesa (Cambrien, SW Sardaigne): Biostratigraphie, paléogéographie, paléoécologie des Archéocyathes: Bulletin de la Société Géologique de France, ser. 8, v. 1, no. 4, p. 531–540. Debrenne, F., and Zamarreño, I., 1970, Sur la découverte d’Archéocyathes dans le Cambrien du NW de l’Espagne: Brevoria Geológica Astúrica, v. 14, p. 1–11. Dies, M.E., and Gozalo, R., 2004, Agnostida (Trilobita) de la Formación Valdemiedes (Leoniense: Cámbrico Medio basal) de las Cadenas Ibéricas (NE de España): Boletín Geológico y Minero, v. 115, no. 4, p. 683–698. Dies, M.E., Gozalo, R., and Liñán, E., 2004, Zonación del Límite Bilbiliense– Leoniense (Formación Valdemiedes, Cámbrico Inferior–Medio) en las Cadenas Ibéricas: Geo-Tema, v. 6, no. 2, p. 283–286. Dies Álvarez, M.E., 2004, Bioestratigrafía y paleoecología de la Formación Valdemiedes (límite Cámbrico Inferior–Medio) en las Cadenas Ibéricas [tesis doctoral]: Zaragoza, Spain, Universidad de Zaragoza, 147 p. Elicki, O., 1997, Biostratigraphic data of the German Cambrian–present state of knowledge: Freiberger Forschungsheft, v. C, no. 466, p. 155–165.
372
Gozalo et al.
Elicki, O., and Pillola, G.L., 2004, Cambrian microfauna and palaeoecology of the Campo Pisano Formation at Gutturu Pala (Iglesiente, SW Sardinia, Italy): Bolletino della Società Paleontologica Italiana, v. 43, no. 3, p. 383–401. Elicki, O., Schneider, J., and Shinaq, R., 2002, Prominent facies from the Lower/Middle Cambrian of the Dead Sea area (Jordan) and their palaeogeographical significance: Bulletin de la Société Géologique de France, v. 173, no. 6, p. 547–552, doi: 10.2113/173.6.547. Fletcher, T.P., 2003, Ovatoryctocara granulata: The key to a global Cambrian Stage boundary and the correlation of the Olenellid, Redlichiid and Paradoxidid Realms: Special Papers in Palaeontology, v. 70, p. 73–102. Fletcher, T.P., 2005, Holaspid variation in the solenopleurid trilobite Parasolenopleura gregaria (Billing, 1865) from the Cambrian of Newfoundland: Palaeontology, v. 48, no. 5, p. 1075–1089, doi: 10.1111/j.1475-4983.2005.00494.x. Fletcher, T.P., Theokritoff, G., Lord, G.D., and Zeoli, G., 2005, The early paradoxidid harlani trilobite fauna of Massachusetts and its correlatives in Newfoundland, Morocco, and Spain: Journal of Paleontology, v. 79, no. 2, p. 312–336, doi: 10.1666/0022-3360(2005)079<0312:TEPHTF>2.0.CO;2. Gámez, J.A., Fernández-Nieto, C., Gozalo, R., Liñán, E., Mandado, J., and Palacios, T., 1991, Bioestratigrafía y evolución ambiental del Cámbrico de Borobia (Provincia de Soria, Cadena Ibérica Oriental): Cuaderno do Laboratorio Xeolóxico de Laxe, v. 16, p. 251–272. García Hernández, M., and Liñán, E., 1983, Estromatolitos y facies asociadas en la Formación Santo Domingo (Cámbrico inferior de la Sierra de Córdoba), in Comba, J.A., coord., Libro Jubilar J.M. Ríos, v. 3: Madrid, Instituto Geológico y Minero de España, p. 125–132. Geyer, G., 1989, Late Precambrian to early Cambrian lithostratigraphy of southern Morocco: Beringeria, v. 1, p. 115–143. Geyer, G., 1990a, Die marokkanischen Ellipsocephalidae (Trilobita: Redlichiida): Beringeria, v. 3, p. 1–363. Geyer, G., 1990b, Revised Lower to Lower Middle Cambrian biostratigraphy of Morocco: Newsletter on Stratigraphy, v. 22, no. 2–3, p. 53–70. Geyer, G., 1994, Cambrian Corynexochid trilobites from Morocco: Journal of Paleontology, v. 68, no. 6, p. 1306–1320. Geyer, G., 1998, Intercontinental, trilobite-based correlation of the Moroccan early Middle Cambrian: Canadian Journal of Earth Sciences, v. 35, p. 374–401, doi: 10.1139/cjes-35-4-374. Geyer, G., and Landing, E., 1995, The Cambrian of the Moroccan Atlas regions: Beringeria, v. 2, Special Issue, p. 7–46. Geyer, G., and Landing, E., 2001, Middle Cambrian of Avalonian Massachusetts: Stratigraphy and correlation of the Braintree trilobites: Journal of Paleontology, v. 75, p. 116–135, doi: 10.1666/0022-3360(2001)075<0116: MCOAMS>2.0.CO;2. Geyer, G., and Landing, E., 2004, A unified Lower–Middle Cambrian chronostratigraphy for West Gondwana: Acta Geologica Polonica, v. 54, p. 179–218. Geyer, G., and Shergold, J.H., 2000, The quest for internationally recognized divisions of Cambrian time: Episodes, v. 23, p. 188–195. Geyer, G., and Wiefel, H., 1997, Fränkisch-Thüringisches Schiefergebirge: Courier Forschungsinstitut Senckenberg, v. 200, p. 56–102. Geyer, G., Landing, E., and Heldmaier, W., 1995, Faunas and depositional environments of the Cambrian of the Moroccan Atlas regions: Beringeria, v. 2, Special Issue, p. 47–119. Gozalo, R., and Liñán, E., 1995, Leonian (early Middle Cambrian) Paradoxides biostratigraphy: Beringeria, v. 2, Special Issue, p. 169–171. Gozalo, R., and Liñán, E., 1997, Revision of the genus Onaraspis Öpik 1968 and its biostratigraphic and biogeographic significance, in Second International Trilobite Conference, Brock University, St. Catharine’s, Ontario, August 22–25, Abstract with Program, p. 24–25. Gozalo, R., Liñán, E., and Álvaro, J.J., 1993a, Bioestratigrafía del Cámbrico Medio de Villafeliche (prov. Zaragoza, España): Revista Española de Paleontología, nº extraordinario, p. 49–57. Gozalo, R., Álvaro, J., Liñán, E., Sdzuy, K., and Truyols, J., 1993b, La distribución de Paradoxides (Acadoparadoxides) mureroensis Sdzuy, 1958 (Cámbrico Medio basal) y sus implicaciones paleobiogeográficas: Cuadernos do Laboratorio Xeolóxico de Laxe, v. 18, p. 217–230. Gozalo, R., Mayoral, E., Gámez Vintaned, J.A., Dies, M.E., and Muñiz, F., 2003, A new occurrence of the genus Tonkinella in northern Spain and the Middle Cambrian intercontinental correlation: Geologica Acta, v. 1, no. 1, p. 121–126. King, W.B.R., 1923, Cambrian fossils from the Dead Seas: Geological Magazine, v. 60, p. 507–514.
Korovnikov, I.V., 2001, Lower and Middle Cambrian boundary and trilobites from northeast Siberian platform, in Peng, S.C., Babcock, L.E., and Zhu, M., eds., Cambrian System of South China, v. 13: Palaeoworld: Hefei: Press of the University of Science and Technology of China, p. 270–275. Korovnikov, I.V., 2004, Lower and Middle Cambrian boundary in stratotype sections of the Siberian Platform, in IX International Conference of the Cambrian Stage Subdivision Working Group, September, 2004, Taebaek, Korea, p. 12–15. Lin, T.-R., and Peng, S., 2004, New material of Palaeolenus (trilobite, Cambrian) from the eastern Yangtze Gorge area, Western Hubei: Acta Palaeontologica Sinica, v. 43, no. 1, p. 32–42. Liñán Guijarro, E., 1978, Bioestratigrafía de la Sierra de Córdoba: Tesis doctorales de la Universidad de Granada, v. 191, p. 212. Liñán, E., and Gozalo, R., 1986, Trilobites del Cámbrico inferior y medio de Murero (Cordillera Ibérica): Memorias del Museo Paleontológico de la Universidad de Zaragoza, v. 2, p. 1–104. Liñán, E., and Gozalo, R., 2001, Aragotus, nuevo género de trilobites del Cámbrico Inferior y Medio de Aragón, in Gámez Vintaned, J.A., and Liñán, E., eds., La era Paleozoica: El desarrollo de la vida marina: Zaragoza, Spain, Institución “Fernando el Católico,” p. 267–276. Liñán, E., Perejón, A., and Sdzuy, K., 1993a, The Lower–Middle Cambrian stages and stratotypes from the Iberian Peninsula: A revision: Geological Magazine, v. 130, no. 6, p. 817–833. Liñán, E., Fernández-Nieto, C., Gámez, J.A., Gozalo, R., Mayoral, E., MorenoEiris, E., Palacios, T., and Perejón, A., 1993b, Problemática del límite Cámbrico Inferior–Medio en Murero (Cadenas Ibéricas, España): Revista Española de Paleontología, nº extr., p. 26–39. Liñán, E., Álvaro, J., Gozalo, R., Gámez-Vintaned, J.A., and Palacios, T., 1995, El Cámbrico Medio de la Sierra de Córdoba (Ossa-Morena, S de España): Trilobites y paleoicnología: Implicaciones bioestratigráficas y paleoambientales: Revista Española de Paleontología, v. 10, no. 2, p. 219–238. Liñán, E., Sdzuy, K., Alvaro, J.J., Gámez Vintaned, J.A., and Gozalo, R., 1996, The Jarque section: Lower and Middle Cambrian (West Asturan–Leonese Zone, Eastern Iberian Chain), in Liñán, E., Gámez Vintaned, J.A., and Gozalo, R., eds., II Field Conference of the Cambrian Stage Subdivision Working Groups, International Subcommission on Cambrian Stratigraphy, Spain, 13–21 September, Field Trip Guide and Abstracts: Universidad de Zaragoza, p. 57–72. Liñán, E., Gozalo, R., Palacios, T., Gámez Vinaned, J.A., Ugidos, J.M., and Mayoral, E., 2002, Cambrian, in Gibbons, W., and Moreno, T., eds., The Geology of Spain: London, Geological Society, p. 17–29. Liñán, E., Dies, M.E., and Gozalo, R., 2003, A review of the genus Kingaspis (Trilobita, Lower Cambrian) from Spain and its biostratigraphical consequences for correlation in the Mediterranean subprovince: Revista Española de Paleontología, v. 18, p. 2–14. Liñán, E., Perejón, A., Gozalo, R., Moreno-Eiris, E., and Oliveira, J.T. de, 2004a, The Cambrian System in Iberia: Instituto Geológico y Minero, Cuadernos del Museo Geominero, v. 3, p. 1–63. Liñán, E., Dies, M.E., Gámez Vintaned, J.A., Gozalo, R., Mayoral, E., Muñiz, F., and Chirivella, J.B., 2004b, Hallazgo del trilobites Alueva hastata (Sdzuy, 1958) y de pistas fósiles en el Cámbrico Medio temprano de Sierra Morena: Consecuencia regionales y ambientales: Miscelánea en homenaje a Emiliano Aguirre, v. II: Paleontología: Zona Arqueológica, v. 4, p. 256–263. Loi, A., Pillola, G.L., and Leone, F., 1995, The Cambrian and Early Ordovician of south-western Sardinia: Rendiconti del Seminario della Facoltà di Scienze dell’Univeristà di Cagliari, suppl. vol. 65, p. 61–81. Ludwig, V., 1969, Lithologische Untersuchung des Kambriums im Frankenwald (Bayern): Geologisches Jahrbuch, v. 87, p. 89–156. Öpik, A.A., 1975, Cymbric Vale Fauna of New South Wales and Early Cambrian biostratigraphy: Bureau of Mineral Resources of Australia Bulletin, v. 159, p. i–iv, 1–78. Palacios, T., and Moczydlowska, M., 1998, Acritarch biostratigraphy of the Lower–Middle Cambrian boundary in the Iberian Chains, province of Soria, northeastern Spain: Revista Española de Paleontología, nº extr., Homenaje al Prof. Gonzalo Vidal, p. 65–82. Parnes, A., 1971, Late Lower Cambrian trilobites from the Timna area and Har’Amran (Southern Negev, Israel): Israel Journal of Earth Sciences, v. 20, p. 179–205. Peng, S., 2003, Chronostratigraphic subdivision of the Cambrian of China: Geologica Acta, v. 1, no. 1, p. 135–144. Peng, S.C., and Babcock, L.E., 2001, Cambrian of the Hunan-Guizhou region, South China, in Peng, S.C., Babcock, L.E., and Zhu, M., eds. Cambrian
The Lower–Middle Cambrian boundary in the Mediterranean subprovince System of South China, v. 3: Palaeoworld: Hefei, Press of the University of Science and Technology of China, p. 3–51. Peng, S.C., Babcock, L.E., Robison, R.A., Lin, H.L., Rees, M.N., and Saltzman, M.R., 2004, Global Standard Stratotype-section and Point (GSSP) of the Furongian Series and Paibian Stage (Cambrian): Lethaia, v. 37, p. 365– 379, doi: 10.1080/00241160410002081. Perejón, A., 1994, Palaeogeographic and biostratigraphic distribution of Archaeocyatha in Spain: Courier Forschungsinstitut Senckenberg, v. 172, p. 341–354. Perejón, A., and Moreno-Eiris, E., 2003, Arqueociatos del Bilbiliense (Cámbrico Inferior) del manto del Esla, Cordillera Cantábrica, Norte de España: Boletín de la Real Sociedad Española de Historia Natural (sección geológica), v. 98, p. 51–71. Perejón, A., Fröhler, M., Bechstädt, T., Moreno-Eiris, E., and Boni, M., 2000, Archaeocyathan assemblages from the Gonnesa Group, Lower Cambrian (Sardinia, Italy) and their sedimentologic context: Bolletino della Società Paleontologica Italiana, v. 39, p. 257–291. Pillola, G.L., 1991, Trilobites du Cambrien inférieur du SW de la Sardaigne, Italie: Palaeontographica Italica, v. 78, p. 1–174. Pillola, G.L., Leone, F., and Loi, A., 1995, The Lower Cambrian Nebida Group of Sardinia: Rendiconti del Seminario della Facoltà di Scienze dell’Univeristà di Cagliari, suppl. vol. 65, p. 27–60. Rushton, A.W.A., and Powell, J.H., 1998, A review of the stratigraphy and trilobite faunas from the Cambrian Burj Formation in Jordan: Bulletin of the Natural History Museum, Geology Series, v. 54, p. 131–146. Sdzuy, K., 1958, Neue Trilobiten aus dem Mittelkambrium von Spanian: Seckenbergiana lethaea, v. 39, p. 235–253. Sdzuy, K., 1961, Das Kambrium Spaniens, Teil II: Trilobiten: Akademie der Wissenschaften und der Literatur, Abhandlungen der mathematisch-naturwissenschaftlichen Klasse, v. 1961, nos. 7–8, p. 499–690 (217–408). Sdzuy, K., 1968, Trilobites del Cámbrico Medio de Asturias: Trabajos de Geología, Universidad de Oviedo, year 1967, v. 1, p. 77–133. Sdzuy, K., 1969, Bioestratigrafía de la Griotte Cámbrica de los Barrios de Luna (León) y de otras sucesiones comparables: Trabajos de Geología, Universidad de Oviedo, year 1968, v. 2, p. 45–58. Sdzuy, K., 1970, Mittelkambrische Fossilien aus neuen Bohrungen bei Doberlug und bei Delitzsch: Geologie, v. 19, no. 9, p. 1066–1091. Sdzuy, K., 1971a, Acerca de la correlación del Cámbrico inferior de la Península Ibérica: I Congreso Hispano-Luso-Americano de Geología Económica, Sección 1, Geología, v. 2, p. 753–768. Sdzuy, K., 1971b, La subdivisión bioestratigráfica y la correlación del Cámbrico Medio de España: I Congreso Hispano-Luso-Americano de Geología Económica, Sección 1, Geología, v. 2, p. 769–782.
373
Sdzuy, K., 1972, Das Kambrium der acadobaltischen Faunenprovinz: Zentralblatt für Geologie und Paläontologie, Teil II, v. 1972, nos. 1–2, p. 1–91. Sdzuy, K., 1995, Acerca del conocimiento actual del Sistema Cámbrico y del Límite Cámbrico Inferior–Cámbrico Medio, in Gámez Vintaned, J.A., and Liñán, E., eds., La expansión de la vida en el Cámbrico: Libro homenaje al Prof. Klaus Sdzuy: Zaragoza, Institución “Fernando el Católico,” p. 253–263. Sdzuy, K., and Liñán, E., 1993, Rasgos paleogeográficos del Cámbrico Inferior y Medio del norte de España: Cuadernos do Laboratorio Xeolóxico de Laxe, v. 18, p. 189–215. Sdzuy, K., Liñán, E., and Gozalo, R., 1996, The Middle Cambrian Series, in Liñán, E., Gámez Vintaned, J.A., et al., eds., II Field Conference of the Cambrian Stage Subdivision Working Groups: International Subcommission on Cambrian Stratigraphy, Spain, 13–21 September: Field Trip Guide and Abstracts: Universidad de Zaragoza, p. 16–18. Sdzuy, K., Liñán, E., and Gozalo, R., 1999, The Leonian Stage (early Middle Cambrian): A unit for Cambrian correlation in the Mediterranean subprovince: Geological Magazine, v. 136, no. 1, p. 39–48, doi: 10.1017/ S0016756899002241. Soudry, D., and Weissbrod, T., 1995, Morphogenesis and facies relationships of thrombolites and siliciclastic stromatolites in a Cambrian tidal sequence (Elat area, southern Israel): Palaeogeography, Palaeoclimatology, Palaeoecology, v. 114, p. 339–355, doi: 10.1016/0031-0182(94)00087-O. Sundberg, F.A., 2005, The Topazan Stage, a new Laurentian stage (Lincolnian series–“Middle” Cambrian): Journal of Paleontology, v. 79, no. 1, p. 63– 71, doi: 10.1666/0022-3360(2005)079<0063:TTSANL>2.0.CO;2. Sundberg, F.A., and McCollum, L.B., 2003, Early and Mid Cambrian trilobites from the outer-shelf deposits of Nevada and California, USA: Palaeontology, v. 46, no. 5, p. 945–986, doi: 10.1111/1475-4983.00328. Yuan, K.-X., Zhu, M.-Y., Zhang, J.-M., and Van Iten, H., 2001, Biostratigraphy of Archaeocyathan horizons in the Lower Cambrian Fucheng Section, South Shaanxi province: Implications for regional correlations and Archaeocyathan evolution: Acta Palaeontologica Sinica, v. 40, supp., p. 115–129. Zamarreño, I., 1972, Las litofacies carbonatadas del Cámbrico de l a Zona Cantábrica (NW de España) y su distribución paleogeográfica: Trabajos de Geología, Universidad de Oviedo, v. 5, p. 1–118. Zhuravlev, A.Y., 1995, Preliminary suggestions on the global Early Cambrian zonation: Beringeria, v. 2, Special Issue, p. 147–160.
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Geological Society of America Special Paper 423 2007
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 Martin S. Oczlon* Aldridge Minerals Inc., J.J.Astorstr. 49, 69190 Walldorf/Heidelberg, Germany Antoneta Seghedi Geological Institute of Romania, 1 Caransebes Str., 78344 Bucharest 32, Romania Charles W. Carrigan Department of Geology, Olivet Nazarene University, 1 University Ave., Bourbonnais, Illinois 60914, USA
ABSTRACT The Moesian Platform is a crustal block within southern Europe, located beyond the southwestern margin of the East European craton. Along this margin lie terranes that were accreted to Baltica as part of Far Eastern Avalonia during Late Ordovician–Early Devonian time and terranes that already formed part of Cambrian Baltica, displaced as proximal terranes together with Far East Avalonian terranes. The tectonic history and crustal affinity of the Moesian Platform, however, remain poorly understood. A review of available tectonostratigraphic, paleontological, and geochronologic data suggests that the Moesian Platform comprises four distinct terranes, two with Baltican and two with Avalonian affinities. A fifth terrane, North Dobrogea, lies between the Moesian Platform and the East European craton and records Variscan (Carboniferous) accretion. This accretionary record leads to the paradox that the youngest accreted crust (North Dobrogea) lies closest to the craton, whereas the earlier accreted crust and crust derived from the craton itself are now located more externally. A review of terranes along the southwestern margin of the East European craton, between the North Sea and the Black Sea, suggests that a dextral strike-slip dominated the southwestern Baltican margin during Late Ordovician–Early Devonian accretion of Far Eastern Avalonia, much as is the case in western North America today. Variscan indentation of the Bohemian Massif led to escape-displacement of some Caledonian terranes, and strike-slip displacement during the Mesozoic opening of Mediterranean-style oceanic basins led to the current juxtaposition of Moesian terranes, inverted with respect to their accretionary history. Keywords: East Avalonia, Baltica, terrane accretion, Caledonian, Variscan *E-mail:
[email protected]. Oczlon, M.S., Seghedi, A., and Carrigan, C.W., 2007, 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, 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. 375–400, doi: 10.1130/2007.2423(18). For permission to copy, contact
[email protected]. ©2007 Geological Society of America. All rights reserved.
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INTRODUCTION The Moesian Platform represents a continental block extending ~600 kilometers (km) east-west and 250–300 km north-south (Fig. 1), located mainly in the present territories of Romania and Bulgaria. It belongs to a zone of Caledonian terranes along the SW margin of the East European craton between Denmark and the Black Sea (Fig. 1). The term Caledonian is used here to describe any process related to the Late Ordovician–Early Devonian amalgamation of Avalonia or Avalonia-related terranes, Baltica, and Laurentia, forming the Devonian–early Carboniferous supercontinent Laurussia. However, minor terrane displacements with little or no internal deformation should not be ranked as Caledonian. Some of the Caledonian terranes along the southwestern margin of Baltica are attributed to Avalonia with Late Ordovician docking to Baltica, although others represent proximal Baltican terranes broken off and displaced during the collisional process with Avalonia. The provenance of these terranes is subject to ongoing controversy, with opinions ranging from exclusively Avalonia-related to predominantly Baltica-derived (Franke, 1994; Cocks et al., 1997; Cocks and Fortey, 1998; Pharaoh,
1999; Unrug et al., 1999; Belka et al., 2000, 2002; Katzung, 2001; Z˙ elaz´niewicz et al., 2001; Cocks, 2002; Winchester et al., 2002a,b, 2004; Kalvoda et al., 2003; Winchester, 2003; Nawrocki et al., 2004; Oczlon, 2006). A critical point for evaluating terrane provenance along the southwestern margin of the East European craton has been the Late Neoproterozoic development (ca. 650–543 Ma). This period records accretion and active margins in peri-Gondwana environments such as Avalonia (e.g., Murphy et al., 2004). The entire West Baltican margin was passive or rifting during that period (e.g., Greiling et al., 1999, for Scandinavia; Bakun-Czubarow et al., 2002, for East Poland; Belov et al., 1987, for the Pre-Dobrogean Depression as indicated in Figs. 2 and 3), but there is increasing evidence that all other margins of Baltica record deformation and crustal thickening (Fig. 1). This includes the area from North Norway to Northeast European Russia (the Timanian orogen; Roberts and Siedlecka, 2002), the West Urals (Glasmacher et al., 2001), the Turan Platform (Khain, 1994), the Greater Caucasus (Adamia et al., 1997), and the Scythian Platform (Khain, 1994). The Turan Platform in the southeastern corner of the East European craton has been interpreted as a terrane assemblage with Variscan (Carboniferous) accretion to Baltica
Figure 1. Location of the Moesian Platform in a simplified geological map of Europe (after Oczlon, 2006). Gondwana-derived terranes in the displacement zone along the Trans-European Suture Zone (TESZ) were in contact with Baltica by Late Ordovician time. Baltica is used as a paleogeographic term, while East European craton is used to address its current geological features.
Avalonian and Baltican terranes in the Moesian Platform (Garzanti and Gaetani, 2002) or as a dextral strike-slip assembly of arc terranes along the South Eurasian margin with accretion during Permo-Triassic time (Natal’in and Şengör, 2005). However, the Devonian Donets-Dniepr-Karpinsky rift and related structures of southern Baltica continue across the northern Caspian Sea into the Turan Platform (Fig. 1; Zonenshain et al., 1990; Khain, 1994), which must therefore be a pre-Devonian element of Baltica with Late Neoproterozoic accretion at the latest (585 Ma hornblende schist reported in Khain, 1994). An alternative interpretation is that Carboniferous to Triassic arc magmatism and deformation were superimposed and partly revived earlier structures of the Turan Platform. Given their potential original location on the South Baltican margin, terranes with Late Neoproterozoic metamorphic, deformational, or arc-magmatic records along the southwestern margin of the East European craton are not necessarily of Cambrian–Early Ordovician Gondwana origin, nor does the record of Late Neoproterozoic detrital mica and zircon ages imply Gondwana provenance (Winchester et al., 2004). However, despite a similar record at the scale of the entire Late Neoproterozoic (ca. 100 Ma), differences should emerge upon more detailed
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analysis of tectonometamorphic and magmatic processes at a scale of 10 Ma, allowing for assignment of terranes to one or the other paleocontinent or even to discrete parts of their margins. At the current stage of investigation, this requires a much larger database, in particular for the Baltican margins. The Cambrian– Early Ordovician interval is critical, because terranes attached to one or the other continent were separated by widening oceanic basins (widest in the Early Ordovician; cf. Servais et al., 2003). Terrane affinities in the Moesian Platform and related terranes along the southwestern Baltican margin are addressed here with analysis of the Proterozoic and Paleozoic tectonostratigraphic development, facies and faunal characteristics, and available geochronologic data, potentially pointing to Gondwanan versus Baltican provenance. Paleomagnetic data are helpful but at this stage are considered less relevant for the assessment of the latitudinal displacement of terranes without established apparent polar wander paths. The reliability of isolated data is difficult to assess, and such data may lead to far-reaching but potentially incorrect conclusions (cf. the review of Robardet, 2003). This is exacerbated by the discovery of octupole components for the Paleozoic and Proterozoic magnetic field (Kent and Smethurst, 1998; Torsvik
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Figure 2. Major structural units on the Moesian Platform and selected drillholes reaching pre-Mesozoic rocks (after Romanian Geological Survey data). Dobrogea is a geographic term for the area between the Danube and the Black Sea coast, while the North, Central, and South subdivisions are based on geological units. Drillhole names: BA—Bals¸ (a group of many drillholes); BG—Ba˘ ra˘ganu; BL—Berlescu; BR—Brades¸ti; BU—Budes¸ti; BV—Bordei Verde; CA—Ca˘la˘ras¸i; CH—Chilii; CO—Cobadin; CP—Cumpâna; CU—Cucueţi; DG—Da˘lgodelci; FA—Faures¸ti; GM—Gârla Mare; HB—Ha˘bes¸ti; KA—Kardam; LE—Leu; MA—Marash; MG—Mangalia; MH—Mihalich; MI—Mitrofani; MO—Mogos¸es¸ti; OG—Ograzhden; OL—Oporelu; OP—Opris¸or; PZ—Palazu; SI—Siminoc; SM—Smirna; ST—Stra˘jes¸ti; ŢA—Ţa˘nda˘rei; TO—Totleben; TZ— Tuzla; VA—Vaklino; VE—Vetrino; ZV—Zavoaia. Faint box at upper right outlines the area shown in Figure 3.
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Figure 3. Geological sketch map of Dobrogea (after Romanian Geological Survey data; see Fig. 2 for location). The Peceneaga-Camena Fault is recognized as the eastern boundary of the Moesian Platform, where Central Dobrogea is the only area with outcropping Pre-Mesozoic rocks. In contrast to South and Central Dobrogea, North Dobrogea has a Variscan (Carboniferous) accretionary record and represents an entirely different terrane. Fm.—Formation.
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and Van der Voo, 2002). The octupole component most strongly distorts the midlatitudinal field to shallow inclinations, but the strength of this component and its variations through time are unknown for pre–late Carboniferous times, rendering the meaning of such data uncertain. Based on the available data, we conclude that the Moesian Platform contains four distinct terranes, two of Avalonian and two of Baltican origin, juxtaposed during a long history of Paleozoic and Mesozoic strike-slip displacement. GEOLOGICAL SETTING In East Romania, the faulted margin of the Trans-European Suture Zone (TESZ) splits up into three faults (Figs. 1–3), active mainly during Mesozoic time (Seghedi, 1998, and references therein). The central one, referred to as the Peceneaga-Camena fault, represents the eastern margin of the Moesian Platform. An important fault referred to as the Intra-Moesian fault crosses the
Gondwana/Hun units with Variscan (Carboniferous) accretion to Baltica Late Neoproterozoic turbidites (Histria Fm.) Meta-arc tholeiites, paragneiss, and micaschist (696-643 Ma metmorphism, K-Ar)
Moesian Platform in a northwest-southeast direction and serves as the boundary for the division between the West and East Moesian Platform (Fig. 2). Except for Central Dobrogea, where Neoproterozoic rocks are largely at outcrop, pre-Mesozoic rocks are concealed under Meso-Cainozoic cover, and knowledge about the Paleozoic history comes entirely from drillholes. Variscan accretion of the Balkan terrane and associated deformation is recorded on the southern margin of West Moesia (cf. Haydoutov and Yanev, 1997).The next accretionary stage along the South Laurasian margin is Cimmerian, which led to mild Latest Triassic folding of West Moesian sediments in a foreland setting (Tari et al., 1997). The Alpine Carpathian-Balkan orogen surrounding the Moesian Platform follows Mesozoic small oceanic basins created during breakup of Pangea (e.g., Csontos and Vörös, 2004). During this process the Moesian Platform and southerly adjacent terranes were displaced to the south. Late Cretaceous– Cainozoic convergence led to closure of the Mesozoic basins and to collision around the margins of the Moesian Platform.
Avalonian and Baltican terranes in the Moesian Platform North Dobrogea Terrane The North Dobrogea terrane (ND) has no Caledonian accretionary record, but forms an essential element for the discussion in this article and is therefore described shortly. Critical for its assessment are the Variscan (Carboniferous) deformation (e.g., Seghedi and Oiae, 1995) and the Ordovician–Devonian sedimentary record, principally found in the Tulcea and Maˇcin Units (Figs. 3 and 4). The Tulcea Unit contains Upper Ordovician– Lower Devonian deep-water sediments of slope and deep basinal origin (only age range proven; cf. Seghedi and Oiae, 1995, and references therein) composed of radiolarites, shales, pelagic carbonates, bedded iron carbonates, and distal turbidites. The Maˇcin Unit contains largely shelf sediments, with a good shallow-water faunal record for the Early Devonian (Iordan, 1988). The Tulcea deep-water sediments have no equivalent in the West Baltican Paleozoic, but the entire succession is almost identical to the outer shelf to basinal sediments on the Balkan terrane of West Bulgaria (Fig. 2; cf. Yanev, 1992). The Balkan terrane is clearly of peri-Gondwana provenance as shown by Middle Ordovician trilobites (Gutiérrez-Marco et al., 2003), and the same provenance is therefore probable for the ND. This is corroborated by the SSW-NNE arrangement of the shelf sediments (Maˇcin Unit) and slope or basinal sediments (Tulcea Units), which requires a north- to northeast-facing continental margin, which is in accord with North Gondwana provenance but incompatible with the southwestern Baltican margin. Irrespective of its provenance, the ND records severe Carboniferous–Early Permian deformation and metamorphism (Seghedi et al., 1999, 2004b) at the margin of Southwest Baltica, which has not been recorded in Central and South Dobrogea as discussed later. In Figure 9, found later in this chapter, this overprinting is addressed as transpressional displacement of the ND along the southern Baltican margin. At the continental scale, both the ND and the Balkan terrane form part of the Hun superterrane (Stampfli, 2000), representing a stretch of continental crust that rifted off the eastern part of North Gondwana during Late Ordovician–Silurian time to collide with the southern margin of Laurussia in the late early Carboniferous.
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(biotite), 202 Ma (muscovite), and 193 Ma (muscovite; Kräutner et al., 1988), suggesting resetting at about the Triassic–Jurassic boundary. These rocks possibly form the basement to up to 5000 meters (m) of turbidites and related basinal sediments with Late Neoproterozoic–Early Cambrian acritarchs and Ediacara-type medusoids (Histria Fm. in Fig. 3; Seghedi and Oaie, 1994; Oaie, 1998). A strong detrital volcanic component indicates that the turbidites were sourced from a southerly located active magmatic arc (Oaie et al., 2004), and an overall upward coarsening facies association suggests northward progradation of the source area (Seghedi and Oaie, 1999). Deformation took place in front of a north-verging thrust wedge, leading to upright north-vergent folds with a slightly fanned subvertical axial-planar slaty cleavage. The two oldest K-Ar dates for very low- to low-grade metamorphism are 573 and 572 Ma (whole rock on slate), but ten further whole-rock K-Ar ages for phyllites and slates are in the range of 532–314 Ma (Kräutner et al., 1988). Data on detrital zircons from the Upper Neoproterozoic–Lower Cambrian turbidites yield U-Pb sensitive high-resolution ion microprobe (SHRIMP) ages of 1497 ± 8 Ma and 1050 ± 10 Ma, 603 ± 5 Ma, and 579 ± 7 Ma (Z˙ elaz´niewicz et al., 2001, concordant ages). Lower Paleozoic rocks in Central Dobrogea are found exclusively in drillholes. At Bordei Verde (BV in Fig. 2), the deformed Histria Formation is overstepped by just slightly inclined Lower Ordovician sandstones. The ca. 120-m thick sandstone-shale succession in the Berlescu and Bordei Verdei drillholes (BL and BV in Fig. 2) is reported to contain Arenigian graptolitic and shelly fauna (Rickards and Iordan, 1975, p. 244), followed by 90 m of shales with graptolites of Late Ordovician and Silurian age (Murgeanu and Patrulius, 1963; Iordan and Spasov, 1989). Beju (1972) reports the presence of Late Ordovician acritarchs in black shales (BV), and Early Wenlockian graptolites were determined upsection (Iordan, 1981, 1990). In the neighboring Berlescu drillhole, graptolites indicate the presence of Upper Wenlockian black shales (94 m), which apparently unconformably overlie Upper Ordovician shales (Rickards and Iordan, 1975). The Upper Wenlockian shales are unconformably overlain by 47 m of Lower Devonian sandstones.
Central Dobrogea Palazu Area Central Dobrogea is the only Moesian terrane with outcropping pre-Mesozoic rocks, and its state of investigation is therefore comparatively good. It contains amphibolite-facies metatholeiites and metasediments with economic, probably Kuroko-type, copperzinc massive sulphide mineralization (Wagner, 1977) occurring near the strike-slip boundary to North Dobrogea (Fig. 3). The metatholeiites are geochemically related to an active margin environment (Crowley et al., 2000), as suggested by the mineralization style. The two oldest K-Ar dates for mica schists of 696 Ma (biotite) and 643 Ma (muscovite; Kräutner et al., 1988, and references therein; no error ranges given) may come close to the age of prograde metamorphism. An amphibolite yielded 526 Ma (hornblende), and three further samples of mica schist yielded 228 Ma
All data obtained on pre-Mesozoic rocks from the Palazu area have been derived from drillholes. The presence of two very strong magnetic anomalies near the Black Sea coast (Fig. 3) sparked exploration for iron ore in that area. Drilling revealed the presence of uneconomic banded iron formation and gneisses in medium- to high-grade metamorphic rocks (PZ in Fig. 2). K-Ar radiometric dating of granitoid gneisses yielded ages of 1.78 Ga (microcl.) and 1.71–1.62 Ga (muscovite; Kräutner et al., 1988, and references therein; no error ranges given). However, one mica schist sample yielded a K-Ar whole-rock age of 867 Ma, and two granitoid gneiss samples yielded ages of 729 and 644 Ma (microcline).
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Figure 4. Composite Paleozoic tectonostratigraphic record of Moesian terranes and terranes with more complete stratigraphies, discussed in this article, compared to idealized sections for Southwest Baltica (Denmark, West Baltic Sea, East Poland, West Ukrainian Shield) (Barbu et al., 1970; Brinkmann and Krömmelbein, 1984; Modlinski et al., 1994) and East Avalonia (Anglo-Welsh Region) (Woodcock, 1990; Strachan et al., 1996). North Dobrogea basement data from Seghedi et al. (1999, 2004b) and Savu and Stoian (1998). Letter L, M, or U next to series names means Lower, Middle, or Upper; Middle Ordovician refers to Llanvirnian or Llandeilian; Middle Silurian refers to Wenlockian. B—Baltica-related fauna and facies; G—Gondwanarelated fauna and facies; L—South Laurussia–related fauna and facies.
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Avalonian and Baltican terranes in the Moesian Platform This basement is overlain by a probably Upper Neoproterozoic volcano-sedimentary succession referred to as the Cocos¸u Formation (PZ, BG, CP, and SI in Fig. 2). In its widest intersection, it starts with ~470 m of basalts and basaltic volcaniclastic rocks overlain by ~450 m of green, red, and purple shales with green conglomerates in the upper part. The succession is always truncated at the top by Jurassic limestones. Basalts have a within-plate, alkaline geochemistry (Seghedi et al., 2000). The Cocos¸u Formation is deformed and exhibits very low-grade metamorphism; it has been K-Ar dated with a single slate sample to 547 Ma (whole rock; Kräutner et al., 1988, and references therein). Two drillholes near Tuzla (TZ in Fig. 2) ended in Lower Ludlovian shales overlain by Cretaceous sediments. The Silurian sediments possibly represent part of the Paleozoic cover to the Palazu Proterozoic, but the contact has never been drilled. South Dobrogea All data obtained on pre-Mesozoic rocks from South Dobrogea have been derived from drillholes. The oldest known rocks in South Dobrogea were drilled at Ţa˘nda˘rei (ŢA in Fig. 2) and consist of 226 m of microconglomerates, sandstones, and shales that exhibit very low-grade metamorphism and are lithologically similar to the Histria Formation of Central Dobrogea. Palynological sampling yielded acritarchs assignable to a Neoproterozoic–Early Cambrian age range (Iliescu in Iordan, 1971). Above a tectonized contact follow 413 m of mature sandstones and shales with minute acrotretid brachiopods and acritarchs dated as Tremadocian (Iliescu and Beju in Iordan, 1971). The ensuing 42 m of shales and tuffites contain Arenigian graptolites and microfossils (Iliescu in Iordan, 1971). The tuffites contain quartz, mica, feldspar, and silicified tuffs, suggesting acidic, probably dacitic, volcanism. Beju (1972) reports Caradocian–Ashgillian acritarchs from an unspecified level in the Ţa˘nda˘rei drillhole, and it is possible that the Middle(?)–Upper Ordovician is highly condensed or reworked. The Ordovician tuffites and shales are overlain by Lower Wenlockian to Prˇidolian black shales and carbonates with an angular unconformity of ~15°. Two volcaniclastic sandstone layers are known in Ludlovian sediments; one consists of angular clasts of feldspars, quartz, amorphous matrix, and sparry calcite cement, the other of crystaloclasts of quartz, feldspar, and biotite and amorphous siliceous vitroclasts within a siliceous, vitroclastic matrix. Lower Ordovician mature sandstones were also intersected in the Mangalia area (MG and CO in Fig. 2). Paraschiv et al. (1983a) report the following acritarch association (ŢA and CO holes): Archaeohystricosphaeridium arenigium, A. minor, Cymatiosphaera boulardii, Cymatiogalea polygonomorpha, Cymatiogalea bellicosa, and Leiofusa species (sp.), indicating a Lower Ordovician, possibly Upper Cambrian, age. The most significant of these are Cymatiogalea bellicosa and Cymatiosphaera boulardii, which are exclusively related to peri-Gondwanan cold-water environments and do not occur on the temperate Early Ordovician shelves of Baltica (Servais, personal commun.). Lower Ludlovian
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graptolite shales with occasional carbonate intercalations overstep lower Ordovician sandstones at Mangalia. At Ca˘la˘ras¸i (CA in Fig. 2), the Lower Ordovician is paraconformably overlain by Lower Ludlovian shales, but an angular unconformity of ~30° is present between the Lower and Upper Ludlovian (Ra˘ileanu et al., 1967). Upper Ludlovian and younger sediments are flat-lying at Ca˘la˘ras¸i, and there is a conformable transition to ~450 m of Lower Devonian–Eifelian sandstones and shales. Lochkovian chitinozoans with North Gondwanan affinities were extracted from various drillholes (CA, ŢA, and ZV in Fig. 2; Vaida et al., 2005). Although there is a conformable Silurian–Devonian transition in western South Dobrogea (ZV and CA in Fig. 2), Lochkovian shales overlie Lower Ludlovian shales with an angular unconformity of 38° at Mangalia (MG; Ra˘ileanu et al., 1967). Some 300–600 m of fossil-dated Lower Devonian–Eifelian sandstones and shales with Rhenish fauna and placoderm fish remains were intersected in various drillholes (SM, CA, and MG in Fig. 2). Land plants are particularly abundant in the Eifelian, and the purple to reddish colors are reminiscent of the Old Red Sandstone facies (Paraschiv et al., 1983a). The clastic succession is overlain by up to 2500 m of Givetian–lower Carboniferous carbonates (CA in Fig. 2; Ra˘ileanu et al., 1967). West Moesian Platform All data obtained on pre-Mesozoic rocks of the West Moesian Platform have been derived from wells, drilled mainly for oil. The metamorphic basement has been reached in an area referred to as the Craiova-Bals¸-Optas¸i basement uplift in the northern West Moesian Platform (see respective drillholes in Fig. 2). Metamorphic rocks were encountered as amphibole schist (566 ± 11 Ma, K-Ar, muscovite, OL in Fig. 2), mica schist (543 ± 17 Ma, K-Ar, mixed biotite + hornblende, BU in Fig. 2; 549 ± 16 Ma, muscovite, ST in Fig. 2), and paragneiss (all data from Paraschiv et al., 1982, 1983b). The ages are interpreted as dating greenschist-facies retrogression (Savu and Paraschiv, 1982). Upper Silurian and Jurassic sediments overstep the metamorphic rocks (Iordan, 1984; Tari et al., 1997) after periods of immediately prior uplift. The oldest nonmetamorphic rocks are >250 m of Lower Cambrian arkosic sandstones and quartzites overlain by 150 m of black shales with occasional carbonate beds. The lower part of the black shale succession contains an early Middle Cambrian trilobite association with closest affinities to Scandinavia and northern Poland (Ha˘bes¸ti and Mitrofani drillholes, HB and MI in Fig. 2; Mutiu, 1991), notwithstanding that some species also occur in peri–North Gondwanan terranes. The Baltican affinities of the trilobite association were confirmed on later reexamination (Rushton and McKerrow, 2000), in accord with a succession of Cambrian–Lower Ordovician facies as recorded similarly in West Baltica (Fig. 4). The early Middle Cambrian at Mitrofani is overlain by ~100 m of Middle and Upper Cambrian shales and black shales dated with acritarchs (Olaru, 1999), and the occurrence of one chitinozoan genus (Lagenochitina) suggests extension into
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the Lower Ordovician. This succession is paraconformably (?) overlain by Wenlockian–Lower Ludlovian shales and tuffites (Paraschiv, 1986; Muţiu, 1991). In drillhole Bals¸ 28 (BA in Fig. 2), a 200-m succession of shales and black shales contains Early Ordovician acritarchs (BA in Fig. 2; Beju, 1972) and overlies undated, but probably Lower Cambrian, red-purple and gray sandstones and shales with an angular unconformity of ~10° (Paraschiv, 1974). Middle–Upper Ordovician sediments appear to be absent on the West Moesian Platform. On the southern West Moesian Platform, 113 m of undated gray-black schists form the oldest known deposits (Middle Cambrian–Lower Ordovician?), intersected at dips ranging from 54° to 90° (Vetrino drillhole, VE in Fig. 2; Yanev and Boncheva, 1995). These rocks are unconformably overlain by deformed shales with occasional carbonate beds of graptolite-dated Llandoverian age (Monograptus sedgwicki) and conodont-dated Wenlockian age (Spassov et al., 1978) ~100 m thick. Llandoverian shales and carbonates are elsewhere reported only from the very western part of the West Moesian Platform (Gârla Mare, GM in Fig. 2; Iordan, 1984). Wenlockian shales and carbonates are known from three locations (GM, CU, and MI in Fig. 2), but elsewhere in the north, Lower and Middle Silurian sediments were apparently not deposited or were eroded before Late Silurian time. Lower Ludlovian shales with intercalated carbonates and occasional tuffites have a wide distribution on the West Moesian Platform, transgressing in the north on metamorphic basement and Middle Cambrian–Ordovician black shales (Paraschiv, 1974). However, at Cucueţi, Gârla Mare, and Vetrino (CU, GM, and VE in Fig. 2), Ludlovian deposits overlie Wenlockian shales conformably (Iordan, 1981; Iordan et al., 1985). Although spanning just a small time interval (6–7 Ma), Upper Silurian shales and carbonates have a true thickness of 300– 900 m on the southern and western West Moesian Platform (GM, OP, and VE in Fig. 2). Although in the south the record includes a complete Ludlovian–Prˇ idolian succession, in the north the record comprises only the Lower Ludlovian (Iordan, 1984), with thicknesses of 115–175 m. Both the thickness and the completeness of the Upper Silurian increase from north to south on the West Moesian Platform, and volcanism appears to be more common in the north. In the southern and western West Moesian Platform, several hundred meters of Lower Devonian shales with occasional sandstone beds conformably overlie the Upper Silurian shales and carbonates (GM, OP, DG, VE, and MH in Fig. 2; Spassov et al., 1978; Iordan, 1984; Lakova, 1995). Much as in South Dobrogea, Lochkovian chitinozoans from various drillholes (DG, VE, and MH in Fig. 2; Lakova, 1995) show affinities to North Gondwana, but the contemporaneous land plant–derived miospore assemblages recovered from two drillholes (DG and KA in Fig. 2) is clearly Laurussia-related (Steemans and Lakova, 2004). On the northern West Moesian Platform, 90–875 m of Lower Devonian shales and sandstones unconformably overlie lower Paleozoic sediments (Murgeanu and Patrulius, 1963; Para-
schiv, 1974; CH, FA, and BA in Fig. 2). Middle–Upper Devonian carbonates were deposited over most of the West Moesian Platform (Vinogradov and Popescu, 1984). However, the carbonate and partly evaporitic platform development starts by Eifelian time in the southern West Moesian Platform (Yanev and Boncheva, 1997), earlier than in the northern West Moesian Platform and the East Moesian Platform (Givetian; Vinogradov and Popescu, 1984). A basal Middle Devonian limestone conglomerate is recorded on the southern West Moesian Platform, locally with slight angular unconformity over Lower Devonian beds (Yanev and Boncheva, 1995, 1997; drillholes TO, MA, VA, OG, and KA in Fig. 2). Intrusives were encountered in two drillholes, with K-Ar ages of 371 ± 11 (MO, biotite in granite; Fig. 2) and 350 ± 11 Ma (LE, hornblende in diorite; Paraschiv et al., 1982, 1983b). Perhaps related Upper Viséan tuffites are recorded at Bra˘des¸ti (BR in Fig. 2; Paraschiv, 1986). Six dioritic intrusives and five basaltic volcanics from various drillholes of the West Moesian Platform yielded exclusively Permo-Triassic K-Ar ages (289–220 Ma; Paraschiv et al., 1982, 1983b). TERRANE DEFINITION The Moesian Platform contains areas with different metamorphic, stratigraphic, and faunal properties, which are here discussed in terms of terrane analysis. For the purpose of this discussion, we use the following definition (Keppie, 1989, with modifications): A terrane is characterized by internal continuity of geology, including stratigraphy, fauna, flora, structure, age and style of metamorphism and deformation, igneous rocks, metallogeny, the geophysical record, and the time of docking or along-margin displacement. It is bounded on all sides by faults or thrusts that may contain trench complexes such as mélanges and ophiolites; it may consist only of a thrust sheet without crustal roots. An exotic terrane drifted across an ocean and contains a geological record different from that of the area it is accreted to, not explicable by facies changes. Examples are microcontinents or island arcs. A proximal terrane is broken off and laterally displaced along the same margin, and there is a significant bounding strike-slip zone, which may be obliterated by later collision (concealed by thrust units, inverted sense of displacement, partial transformation into a reverse fault or thrust, and syncollisional rotation). A proximal terrane may contain a geological record similar to the area against which it is juxtaposed.
TERRANE PROVENANCE The geological aspects of each terrane contained in the Moesian Platform allow for some deductions on terrane provenance, but the largely drillhole-based record remains necessarily incomplete and often allows for more than one interpretation. The earlier view of the Moesian Platform was that it has been a geological entity since Precambrian times, but some traits of the individual
Avalonian and Baltican terranes in the Moesian Platform units (the West Moesian Platform, Central and South Dobrogea, and the Palazu area) reveal incompatibilities, suggesting that they represent terranes under the definition given earlier. The authors recognize that this view is preliminary and must be subject to further investigations on available drilled core material. Central Dobrogea Terrane (CD) The oldest K-Ar data of 696 and 643 Ma on amphibolitefacies rocks in the CD may come close to the terrane’s metamorphic age. These data are comparable with the evolution of the British Midlands (Eastern Avalonia), where a ca. 650 Ma lower amphibolite-facies metamorphic event is recorded (Strachan et al., 1996). Notable among the pre–Late Neoproterozoic detrital zircons is the presence of a Grenvillian (1.05 Ga) and “Rondonian” (1.497 Ga) record and the apparent absence of a Paleoproterozoic source, striking even at the low number of dated grains reported by Z˙ elaz´niewicz et al. (2001). A Grenvillian ( = Sveconorwegian) record exists on the western margin of Baltica in southwestern Scandinavia (Fig. 1), where granitoids with ages of 1.05 and 1.00 Ga have been reported (e.g., summary in Bingen et al., 2003). Nevertheless, the lack of any Late Neoproterozoic collisional or accretionary record on the western Baltican margin precludes provenance of the CD from that area. An Amazonian source for the CD is compatible with a Grenvillian and Rondonian detrital record (cf. Winchester et al., 2002a for the significance of ca. 1.5 Ga “Rondonian” or Amazonia-related sources). Although there is a strong detrital volcanic component in the Upper Neoproterozoic–Lower Cambrian turbidites, lack of any igneous activity in the sedimentary basin argues against a back-arc or intra-arc setting, but indicates a forearc environment with a southerly located arc for the dominant northerly transport directions. This setting is not plausible for the southern and southwestern Baltican margins, but is consistent with a location on the North Gondwana margin as suggested by Oaie et al. (2004). The data suggest that the CD is an Avalonia-related terrane rather than one of Baltican provenance before Ordovician time. Clearly more detrital zircon data are needed from the Upper Neoproterozoic–Lower Cambrian of Central Dobrogea, but also from the overstepping Lower Ordovician sandstones in the Bordei Verde drillhole (BV in Fig. 2). Palazu Terrane (PZ) Rocks with Late Paleoproterozoic metamorphic ages, as found in the PZ (1.78–1.62 Ga) are common on the southwestern East European craton (e.g., Bogdanova et al., 2001). In contrast, Avalonian basement typically records crust-forming and accretionary events at ca. 1.2–0.9 Ga and 0.7–0.55 Ga (Keppie et al., 2003; Murphy et al., 2004). Assignment of the PZ to Baltica is corroborated by the within-plate, alkaline nature of the Upper Neoproterozoic Cocos¸u volcano-sedimentary succession,
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a minimum of ~900 m thick, which overlies the Upper Paleoproterozoic metamorphic rocks. Such successions were widely deposited on the rifted western margin of Baltica during the Late Neoproterozoic (Compston et al., 1995; Greiling et al., 1999; Bakun-Czubarow et al., 2002), suggesting a similar setting for the Cocos¸u Formation. In its presumed original location on the southwestern corner of Baltica, the Cocos¸u rift succession was located in the transition zone of the West Baltican rifts to the collisional or transpressional margin of South Baltica, as indicated in 600–550 Ma reconstructions of Keppie et al. (2003) or Murphy et al. (2004). In this environment, the Cocos¸u rift succession may have been inverted during transpressional terrane juxtaposition on the South Baltican margin, leading to deformation and metamorphism dated to ca. 547 Ma (Kräutner et al., 1988). This overprint may have led to the observed partial resetting of K-Ar ages in Upper Paleoproterozoic medium-grade metamorphic rocks. South Dobrogea Terrane (SD) The geological features of the SD are very similar to those of the CD (Fig. 4), and both terranes appear related. The record of Early Ordovician acritarchs diagnostic for a peri–North Gondwana cold-water environment underlines the Avalonia-related derivation of the SD. The shelly faunal association in Upper Silurian carbonates is very similar to the association on the western East European craton (Iordan, 1999), and the Devonian facies, fauna, and flora show a typical South Laurussian development (Old Red facies, Middle Devonian–lower Carboniferous carbonate platform). As discussed in the following section, this does not contradict finds of Lowermost Devonian chitinozoans with North Gondwanan affinities. Once probably continuous with the CD, the SD was broken off and juxtaposed next to the CD during Mesozoic time (cf. Fig. 5). West Moesia Terrane (WM) The key argument for a Baltican provenance of the WM is the presence of an early Middle Cambrian trilobite association diagnostic for West Baltica (Muţiu, 1991; Rushton and McKerrow, 2000). This is in accord with a typical West Baltican succession of Lower Cambrian sandstones and conglomerates, Middle Cambrian black shales with occasional carbonate intercalations, and Upper Cambrian–Lower Ordovician black shales (Fig. 4). However, the basement of the northern WM records Latest Neoproterozoic metamorphism (566–543 Ma), which cannot be reconciled with the contemporaneous rift development along West Baltica, but rather suggests derivation from the margin of South Baltica, perhaps south of the western Scythian Platform (Figs. 1 and 6). Studies of detrital zircons in Lower Cambrian sandstones from West Moesia may help to further elucidate the provenance question. Shelly fauna in Upper Silurian carbonates is very similar to the faunal association on the western East European craton
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Figure 5. Caledonian terranes on the southwestern margin of the East European craton. Proterozoic drillhole data after Marheine and ValverdeVaquero (2002) for Warszawa and Mînzatu et al. (1975) for the East Carpathian foredeep; Denmark 800–880 Ma basement after Frost et al. (1981) and Katzung (2001); 1.46–1.36 Ga granites in the Southwest Baltic Sea area after Åhäll et al. (1997), Tschernoster (2000), Cˇ ecˇys et al. (2002), and Obst et al. (2004); 1.56 Ga Rapakivi intrusive after Morgan et al. (2000). BA—Banat (Alpine magmatic arc); CD—Central Dobrogea; CE—Ceahla˘u ophiolites (Alpine); CLP—Carpathian lower Paleozoic; FEAAC—Far East Avalonian accretionary complex (thrust sheet on the Baltican crust); GEM—Gemerides (considered part of the CLP); HEE—Holstein–East Elbe; IZ—Istanbul-Zonguldak; ŁG—Łysogóry; LZW— Lubliniec-Zawiercie-Wielun´; MPK—Małopolska; MV—Moravia; ND—North Dobrogea (Variscan); NSB—North Sea basement; NV—Novaci meta-granite (part of the CLP); PZ—Palazu; RZ—Rzeszotary; SD—South Dobrogea; SJY—South Jylland; SV—Severin ophiolites (Alpine); USL—Upper Silesia; VL—Vlasina Complex (considered part of CLP); WMP—West Moesian Platform. Letters A, C, and V next to transcurrent faults are indicative of Alpine, Caledonian, and Variscan times of displacement.
Avalonian and Baltican terranes in the Moesian Platform
Latest Caradocian (450 Ma)
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Figure 6. Late Ordovician paleogeography during the docking of Far East Avalonia to Southwest Baltica. The Far East Avalonian accretionary complex (FEAAC) was ~ 300 km southwest of its current position on initial docking (455–450 Ma). At an emplacement rate of 2 cm/yr, it could have reached its final position at ca. 440–435 Ma (Late Llandoverian or Early Wenlockian). Terrane patterns and abbreviations as in Figure 5.
(Iordan, 1999), and the Devonian facies, fauna, and flora show a typical South Laurussian development (Old Red facies, Middle Devonian–lower Carboniferous carbonate platform). Lowermost Devonian chitinozoans show North Gondwanan affinities (Vaida et al., 2005), but this relationship may reflect the exchange of planktonic fauna via ocean currents across the Rheic Ocean, potentially spreading chitinozoans to continental margins in similar climatic zones (Paris, personal commun.). Chitinozoans are viewed as floating eggs of a marine, soft-bodied animal (Paris and Nolvak, 1999), and single species are likely to live anywhere within a global climatic belt, to which North Gondwana and South Baltica may have pertained. This is underlined by the presence of North Gondwana–related chitinozoans of the same age in Podolia (western Ukraine), an area definitely pertaining to the Early Devonian shelf sea of Southwest Baltica (Paris and Grahn, 1996). Clear South Laurussian palaeogeographic affinity determined with Lochkovian miospore assemblages (Steemans and Lakova, 2004) is given priority here. This evidence is considered more solid than the presence of chitinozoans, because the according flora belongs to the inner continental margin, and derived spores should better reflect its paleogeographical affinities. The development of Lower Silurian–Lower Devonian facies on the WM indicates that its southern and western margins have a thicker, more complete, and more distal record than the northern WM. Clearly the southern and western margins were grading
into a basinal area throughout the Silurian–Devonian (Fig. 7), and this basin was not involved in collision and uplift during that time (otherwise, a source of detritus would be located around the southern and western margins of the WM). Thus the West Moesia terrane occupied an external position during Caledonian terrane displacement along the southwestern margin of Baltica, facing an open, possibly oceanic area to the southwest. The pattern of unconformities and sedimentation in the northern WM can be interpreted in terms of strike-slip terrane displacement: periods of transpressive displacement led to uplift and erosion, while transtensional displacement led to subsidence. However, the lack of associated deformation suggests only a mild transpressive component. During the period with strongest subsidence, up to 900 m of Upper Silurian sediments were deposited on the WM. Upper Silurian thicknesses of 600–700 m in the same facies are recorded in South Dobrogea (drillholes CA, ŢA, and ZV in Fig. 2), except where erosion at about the Silurian–Devonian boundary removed some of the Upper Silurian sediments (MG in Fig. 2). An unconformity above the Lower Ludlovian succession is recorded on the northern WM and in the adjacent SD, but no Upper Silurian sediments are known in Central Dobrogea. This pattern of Upper Silurian sedimentation suggests that the WM and SD behaved much like a coherent crustal block, with only minor displacements along the terrane boundaries from the Late Silurian onward. Volcanic centers on
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Late Llandoverian (435 Ma)
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Figure 7. Early Silurian paleogeography. Sinistral translation of the Far East Avalonian terranes along the Southwest Baltican margin. Closure of the remnant Thor oceanic basin in the (present) North Sea area. Terrane patterns and abbreviations as in Figure 5.
Early Frasnian (380 Ma) CALEDONIAN FRONT
ac io et cr
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EAST AVA LONIA
SJY LG LZW RZ MPK SD US PZ L MV CD CL WM IZ P HEE
RHEIC OCEAN
Figure 8. Late Devonian paleogeography. Avalonia, Laurentia, and Baltica are united in the supercontinent Laurussia, whose southern margin records the development of rifted arc volcano-sedimentary basins (Rhenohercynian zone s.l.). Note the counterclockwise rotation of East Avalonia during Silurian–Early Devonian time (according to Verniers et al., 2002). Terrane patterns and abbreviations as in Figure 5.
Avalonian and Baltican terranes in the Moesian Platform
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as postulated for the South Laurussian margin during Devonian– early Carboniferous time (Fig. 8). A geochemical supra–subduction zone signature was determined for Permian basic–intermediate volcanic rocks in the northern WM, dated 280–240 Ma (Savu and Paraschiv, 1982), possibly caused by extension-related melting of subduction-modified mantle or by ongoing subduction of oceanic crust (Fig. 9).
transtensional faults can be deduced from occasional tuffite intercalations. The relatively high Late Silurian sedimentation rate of ~10–15 cm/1000 yr indicates that the clays must have been rapidly eroded and transported from a Caledonian uplifted zone. This source must be sought further to the northwest. The generation of clays may have been aided by erosion of predominantly sedimentary successions and rapid weathering in a tropical climate. None of the granitoids encountered in lower Paleozoic rocks or the metamorphic basement (MO in Fig. 2) of the northern WM yielded pre–Late Devonian intrusive K-Ar ages, although Uppermost Neoproterozoic K-Ar ages are commonly obtained for the metamorphic rocks. This rather precludes a systematic dating error or resetting and suggests that none of these intrusives was generated during Caledonian and earlier orogenic processes. In the absence of any evidence for Carboniferous crustal thickening (northern WM), these granites are unlikely to have been a response to Variscan collisional processes. If correct, the ages of 371 Ma for a granite and 350 Ma for a diorite, and the presence of Viséan tuffites, may indicate generation above a subduction zone,
Mode of Terrane Displacement The mixed occurrence of Avalonia-related and Baltican proximal terranes on the Moesian Platform requires major tectonic processes that may have acted either as strike-slip or as thrusting of terrane-scale units, or both. The Paleozoic–recent sedimentary record on all four Moesian terranes precludes largescale overthrusting at any stage. Some successions are flat-lying or weakly tilted from Lower Ordovician or Upper Silurian levels onward, without any tectonic overprinting. Angular unconformities are recorded, but here again, crustal shortening is largely
Carboniferous - Permian Boundary (292 Ma) WEST SIBERIAKAZAKHSTAN TERRANE ASSEMBLAGE
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RZ SD PZ CD IZ BK
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North Gondwana-derived Variscan terranes: Carboniferous accretion to Laurussia
TETHYS OCEAN
Figure 9. Paleogeography following the Carboniferous accretion of Gondwana-derived terranes to Laurussia. Indentation of the Bohemian Massif caused south- to ESE-directed escape of USL, LZW, MPK, and ŁG and intense Carboniferous deformation near the terrane boundaries. The Early Permian metamorphic ages in ND (Seghedi et al., 1999, 2004b) are interpreted to have been caused by cooling after transpressional crustal thickening on the South Laurussia margin. Patterns and abbreviations as in Figure 5; BK—Balkan terrane.
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absent (except for Variscan deformation restricted to the southern margin of the WM). In the absence of any crustal thickening, the Moesian terranes must have assumed their current juxtaposition exclusively via strike-slip. Faults separating the WM, SD, and PZ (Figs. 2 and 3) do not appear to have had a significant Meso–Cainozoic play, because overlying rocks of that age are largely undisturbed. Both faults are Caledonian strike-slip terrane boundaries overstepped by a carbonate platform that is Middle Devonian, at the latest. Geophysical data east of the Intra-Moesian fault indicate a crustal thickness of 40–45 km for the SD, opposed by a thinner crust for the WM (30–33 km, Mocanu et al., 1991), underlining the importance of this fault as a Caledonian terrane boundary. Significant Meso-Cainozoic displacement can be demonstrated for the Capidava-Ovidiu and Peceneaga-Camena fault zones (Fig. 3; see later discussion).
Grenvillian Record on the Southwestern Baltican Margin? To assess terranes via their detrital zircon content, the Grenvillian record (1.2–0.9 Ga) is a critical component. Grenvillian tectonometamorphic and magmatic processes are widely known in Southwest Scandinavia, where they are termed “Sveconorwegian” (e.g., Bingen et al., 2003; Fig. 1). The southwestern Baltican margin, here regarded as the area from the Southwest Baltic Sea to the western margin of the Ukrainian Shield (Fig. 1), has no such record (Fig. 5 and references in figure caption). A simple prolongation of the Southwest Scandinavian Grenvillian structures cannot be expected along Southwest Baltica, because these structures run at an angle of ~90° to the southwestern Baltican margin (Fig. 1) and continue into a long drifted-off Grenvillian zone, perhaps now located around the Amazonian craton (cf. Murphy et al., 2004, for possible location relative to Baltica). Instead the southwestern Baltican margin contains a distinctively pre-Grenvillian record, with zircon-dated, sometimes deformed and metamorphosed granites of 1.46–1.36 Ga (Southwest Baltic Sea area in Fig. 5 and references in figure caption). Ar-Ar, K-Ar, and Rb-Sr ages from the same granites (Obst et al., 2004) and from metamorphic rocks in drillholes farther south along the southwestern Baltican margin (Fig. 5 and references in figure caption) are consistently ca. 1.4– 1.22 Ga. Brewer et al. (2004) deduce active margin processes for the period of 1.46–1.20 Ga in Southwest Scandinavia, and recurring deformation in a strike-slip-dominated active margin setting may have operated all along the southwestern Baltican margin. Hence Upper Neoproterozoic to Cambrian sandstones
PALEOGEOGRAPHIC MODEL The following sections review other Caledonian terranes along Southwest Baltica for comparison with the Moesian terranes in order to provide a framework for paleogeographic reconstructions along the southwestern margin of Baltica (Figs. 6–10). Next the Variscan and Alpine displacement and overprinting are assessed. A review of the principal characteristics of Caledonian terranes along the southwestern margin of the East European craton, between the North Sea and the Black Sea, allows them to be placed in groups outlined in the following paragraphs (for more details, see the references cited in the Introduction section).
Early Campanian (80 Ma)
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Figure 10. Late Cretaceous paleogeography. Caledonian terranes were disrupted and displaced due to opening of Mesozoic oceanic basins. The southern part of the Ceahla˘u-Severin Ocean was subducted to the south, creating the Banat (Southwest Romania)–Timok (eastern Serbia)–Srednogorie (Bulgaria) rifted arc-magmatic belt. Late Cretaceous remnant Vardar Ocean according to Pamic´ et al. (2002). Terrane patterns and abbreviations as in Figure 5.
Avalonian and Baltican terranes in the Moesian Platform deposited on the southwestern Baltican margin, south of the Baltic Sea, may not contain a proximal Grenvillian component (i.e., euhedral 1.2–0.9 Ga zircons), while Grenvillian zircons transported over a long range could occasionally form a component derived from Southwest Scandinavia or Amazonia. Known Southwest Baltican source areas would provide a zircon spectrum with the following ages: Late Archean (3.0–2.5 Ga, Ukrainian Shield area), Middle or Late Paleoproterozoic (2.1–1.8 Ga, ubiquitous; 1.8–1.7 Ga rapakivi granites in Ukrainian Shield), Early Mesoproterozoic (ca. 1.6–1.5 Ga rapakivi granites and 1.46–1.36 Ga granites, South Baltic Sea area), and Late Neoproterozoic (ca. 0.65–0.55 Ga acidic components of rift volcanism, ubiquitous). BALTICAN TERRANES WITH LATE NEOPROTEROZOIC BASEMENT This group comprises terranes with Upper Neoproterozoic granitic and medium- to low-grade metamorphic basement, as found in West Moesia and in the Upper Silesia terrane (USL), Southwest Poland, and the East Czech Republic. Early–Middle Cambrian trilobite associations are more closely related to Baltica than to any other region (Early Cambrian Holmia faunal province). There is a widely distributed, shallow-dipping lower Paleozoic cover that was undeformed during Late Ordovician– Early Devonian time. In the northern USL, the basement is overlain by Uppermost Neoproterozoic, strongly deformed, and low- to very low-grade metamorphosed conglomerates, sandstones, and shales of recycled orogen composition, at least in part turbidites (Z˙ elaz´niewicz et al., 2001). Detrital zircons from these sediments apparently show two clusters of U-Pb SHRIMP-determined ages within 1.971 ± 39 Ga and 693 ± 25 Ma and Late Neoproterozoic ages to 577 ± 3 Ma (Z˙ elaz´niewicz et al., 2001). U-Pb SHRIMPdated zircons of granitoids in the USL yielded an age of 540 Ma (Z˙ elaz´niewicz et al., 2001). K-Ar data for drilled basement rocks in the Southwest USL (East Czech Republic; Dudek and Melková, 1975) suggest intrusive cooling ages of 560–600 Ma for diorites (hornblende and biotite) and metamorphic cooling ages of 610–660 Ma for paragneiss (biotite and muscovite). Lower Cambrian conglomerates and sandstones with Holmia province trilobites (Cocks, 2002) form the overstep sequence, overlain in one drillhole in the northern USL by Middle Cambrian sandstones and shales (Moczydłowska, 1997). The Cambrian beds are paraconformably overlain by Lower Devonian conglomerates and sandstones (Moczydłowska, 1997). If restored to its supposed position before the opening of a Mesozoic oceanic basin in the Carpathian area (Banks and Robinson, 1997), the WM was located immediately south of the USL (Fig. 9). Both terranes probably once formed a single block, detached together from the southern margin of Baltica during Late Ordovician–Silurian time (Fig. 6). This explains the absence of Variscan deformation and the presence of largely flat-lying lower Paleozoic sediments on both the southern
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margin of the USL (as far as recorded in drillholes under the Carpathian foreland; cf. Moczydłowska, 1995) and the northern margin of the WP, although both are currently bordering terranes with strong Variscan deformation and metamorphism (the Carpathians). As further outlined later, the Inner Carpathian region was injected between the USL and the WM after their separation in the Mesozoic. An Early Cambrian location of the USL on the South Baltican margin may explain certain differences between the USLtrilobite association there and that of Scandinavia (cf. Nawrocki et al., 2004). In any case, this trilobite association clearly pertains to the shelves of Baltica and cannot be related to any other paleocontinental area (Cocks, 2002). Crustal properties obtained by Malinowski et al. (2005) for the USL from seismic data reveal significant differences with the Malopolska terrane (MPK) and the western East European craton. This may indicate Late Neoproterozoic modification of the USL crust or Late Neoproterozoic accretion of exotic crust, now underlying the USL, to the southern margin of Baltica. BALTICAN TERRANES WITH LATE NEOPROTEROZOIC RIFT RECORD This group comprises terranes with Late Neoproterozoic volcano-sedimentary rift successions with characteristic green and red-purple colors throughout, originally deposited on the southwestern Baltican margin. Such deposits are present in the Palazu terrane (PZ; Seghedi, own drill core review data), the Małopolska terrane (MPK; Unrug et al., 1999), and the South Jylland terrane (SJY; Franke, 1994; Giese et al., 2001). Possible basement to the Upper Neoproterozoic green-red rift successions is known only in the PZ of Dobrogea, dated 1.78–1.62 Ga (see earlier discussion). South Jylland Terrane (SYJ) The SJY is known only from drillholes where slates have been recovered, but its basement is unknown. Detrital Ar-Ar mica ages of 830 and 890 Ma were obtained on slates (Dallmeyer et al., 1999; Giese et al., 2001) and point to a central West Baltican source area as present, for example, in Denmark (Fig. 5; Katzung, 2001) and southern Norway (Mulch and Cosca, 2004). At the bottom of the Rügen 5 drillhole, 42 m of undated green-red slates were cored below 3130 m of Ordovician shales and distal turbidites of the Far East Avalonian accretionary complex (FEAAC in Fig. 5; Katzung, 2001). These green-red slates are interpreted by Beier and Katzung (2001) as Upper Proterozoic deposits of the West Baltican distal shelf emplaced on an internal region of the Baltican shelf at the base of the FEAAC during Caledonian thrusting (Fig. 6). This also suggests that the green-red slates of the SJY form a thrust unit, but the once-overlying FEAAC was eroded. Today the SJY is separated from the FEAAC by a near-vertical structure (Dallmeyer et al., 1999), probably a Caledonian strike-slip fault.
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Małopolska Terrane (MPK) Similar to the PZ, the Late Neoproterozoic succession of the entirely concealed central and western Małopolska terrane records deformation and very low-grade metamorphism, increasing to the west (Belka et al., 2002). The tilted or deformed Upper Neoproterozoic rocks are overstepped by patchily distributed Lower Ordovician and younger sediments (e.g., Compston et al., 1995, for the central MPK). A tuffite from unmetamorphosed Upper Neoproterozoic deposits in the central MPK yielded a U-Pb SHRIMP zircon age of 549 ± 3 Ma, indistinguishable from the 551 ± 4 Ma age for Southwest Baltican rift volcanism in East Poland (the East European craton; Compston et al., 1995). Similar rift-related sediments and magmatic rocks are reported from the western margin of the MPK (Unrug et al., 1999). The eastern MPK, at outcrop in the Holy Cross Mountains, records a lower Paleozoic succession with fossil-dated Lower Cambrian sandstones and Holmia province trilobites (Cocks, 2002). U-Pb SHRIMP-dated detrital zircons from partly acritarchdated Upper Neoproterozoic sediments in the central and eastern MPK yielded ages of 2748 ± 15 Ma, 2560 ± 15 Ma, 2551 ± 44 Ma, 2128 ± 20 Ma, 2042 ± 85 Ma, 1920 ± 63 Ma, 1171 ± 10 Ma, 611 ± 7 Ma, 560 ± 7 Ma, and 527 ± 15 Ma, and an unspecified cluster was dated at 600–580 Ma (Z˙elaz´niewicz et al., 2001; all ages concordant or concordant within errors). The youngest detrital age, if correct, would indicate a maximum depositional age for the upper part at about the Late Neoproterozoic−Cambrian boundary, in accord with SHRIMP zircon-dated tuffs in the central MPK (cf. Compston et al., 1995). The statistically significant cluster of detrital zircons in the range of 2.75–1.9 Ga, and apparent absence of zircons in the range of 1.35–1.85 Ga (characteristic ages for the East European craton in the area of the Baltic Sea and East Poland) is compatible with provenance from the Ukrainian Shield area (cf. Shchipansky and Bogdanova, 1996) and a more southerly Cambrian location of the MPK at the southwestern Baltican margin. The Early Grenvillian input, if substantiated, may relate to a more distal source in Southwest Scandinavia. Cambrian sediments in the eastern MPK were involved in a deformation event with folding and thrusting at about the Cambrian−Ordovician boundary (Gagała, 2005). The same event is apparently recorded in an unconformity on the WM (Lower Ordovician sediments with angular unconformity over Lower Cambrian sandstones; BA in Fig. 2), the Łysogóry terrane, and the Lubliniec-Zawiercie-Wielun´ terrane (see following sections). Associated tectonic processes on the West Baltican margin are probably reflected in the structural orientation of fossil-dated Upper Cambrian or Lower Ordovician neptunian dikes in Southwest Sweden. These dikes are arranged in strike-slip duplex structures of transpressive and transtensive character, in accord with sinistral displacements on the southwestern margin of Baltica (Greiling et al., 1988). This setting is compatible with the absence of crustal thickening and metamorphism at about the Cambrian–Ordovician boundary in the respective terranes, which is not suggestive of collisional tectonics.
There is a record of Upper Ordovician bentonites on the MKP (Masiak et al., 2003), which correlate with the record in East Laurentia and West Baltica, indicative of vast explosive volcanism at a destructive plate margin in the region (Huff et al., 1992; East Avalonia?). OTHER TERRANES OF BALTICAN PROVENANCE Lubliniec-Zawiercie-Wielun´ Terrane (LZW) The LZW (Unrug et al., 1999) lies between the USL and the MPK (Fig. 5). The oldest recognized rocks are Cambrian siltstones unconformably overlain by Ordovician conglomerates and a thick succession of turbidites with conodont-dated (redeposited?) carbonate intercalations of Arenigian, Llanvirnian, and Llandeilian–Caradocian age. These sediments underwent lowgrade metamorphism in Late Ordovician time and are unconformably overlain by Lower Silurian carbonates. The presence of a major unconformity at about the Cambrian–Ordovician boundary and of Lower Ordovician carbonates is inconsistent with the East Avalonian record but compatible with a location on the outer southwestern Baltican margin near the MPK. The docking of Far Eastern Avalonia to Southwest Baltica probably induced Late Ordovician metamorphism and deformation, followed by overstepping Lower Silurian carbonates. Łysogóry Terrane (ŁG) No rocks older than Middle to Upper Cambrian sandstones up to 1800 m thick are known in the ŁG, but the presence of detrital mica with ages of 1.74–1.72 Ga and 0.77–0.93 Ga (Belka et al., 2000) suggests location on the western Baltican margin by Middle to Late Cambrian time (cf. Dallmeyer et al., 1999, and Giese et al., 2001, for the significance of 830–890 Ma detrital mica). There is a rich Late Cambrian trilobite fauna with mixed Baltican, Avalonian, and South American affinities, which does not allow for a clear paleogeographic assignment (Z˙ylin´ska, 2002). The thick Upper Cambrian sandstones may relate to a transtensional setting on the southwestern Baltican margin next to transpressionally deformed, uplifted, and eroded blocks (e.g., the eastern MPK). Everywhere in the ŁG, Ordovician sediments discordantly overly the Cambrian (Buła and Jachowicz, 1998), and Ordovician beds are in turn disconformably overlain by the Lower Silurian (Unrug et al., 1999). Geophysical data indicate that the ŁG is a narrow, ~30-km wide block separated from the East European craton by a vertical structure (Semenov et al., 1998), suggesting that it may have had some displacement against the southwestern Baltican margin at about the Cambrian–Ordovician boundary (see MPK) and/or after docking of Far Eastern Avalonia in Late Ordovician time. The amount of displacement against the southwestern Baltican margin depends greatly on the southern extent of the Far East Avalonian accretionary complex (FEAAC; cf. Figs. 5 and 6 and the following section), because the ŁG was not overthrust and must therefore
Avalonian and Baltican terranes in the Moesian Platform have been located to the south of the southernmost occurrence of the FEAAC. The fact that the Late Cambrian trilobite fauna, with strong Northwest Gondwanan influence, is different in composition from the (now) neighboring part of the East European craton and Scandinavia (cf. Z˙ylin´ska, 2002) suggests that the ŁG may have occupied a different, perhaps more southerly, position along Southwest Baltica. By Latest Silurian time, sedimentation on the ŁG had the character of a foredeep at the southwestern Baltican margin (cf. Narkiewicz, 2002), with land plants in near-shore sediments (Earliest Prˇ idolian; Bodzioch et al., 2003). Alternatively, thick Late Silurian sediments, basinal in the beginning, may reflect the shallowing of a transtensiongenerated filled-up basin related to the Late Silurian transcurrent regime of the southwestern Baltican margin. Holstein–East Elbe Terrane (HEE)
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Caradocian turbidites, shales, and occasional tuffites with peri– North Gondwanan acritarchs (Servais and Fatka, 1997). Across the Southwest Baltic Sea, the FEAAC recurs in Northwest Poland, and docking has been shown to be pre-Middle Ashgillian for the presence of reworked North Gondwanan microfossils in Southwest Baltican sediments (Samuelsson et al., 2002). Farther south, the FEAAC appears to continue into a zone of strongly deformed, “geosynclinal” Ordovician to (?) Silurian deposits figured by Khizhnjakov (1969) from the area west of Warszawa to SSE of Lublin (drillhole data; Fig. 5). The southern extent of the FEAAC is uncertain, but critical to the amount of displacement of some proximal Southwest Baltican terranes, which cannot have been overthrust during emplacement of the FEAAC, but have a more or less continuous sedimentary record in the Late Ordovician–Early Silurian without deformation in that period (the ŁG and MPK).
No drillhole ever reached the pre-Devonian basement in the deeply buried HEE, but xenoliths in upper Carboniferous volcanic rocks from the western HEE in northern Germany contain anorthosites as found widely on the western margin of Baltica (Kämpf et al., 1994). Anorthosites comparable in position and lithology to the xenoliths from the HEE occur in Southwest Norway and are dated at 0.93 Ga (Rogaland Complex; summary in Bingen et al., 2003). Inherited zircons in upper Carboniferous volcanic rocks from the eastern HEE in northern Germany yielded SHRIMP-dated Pb-Pb ages of 1.48 and 1.46 Ga (Breitkreuz and Kennedy, 1999). Granites of that age occur in Southwest Sweden (cf. Åhäll and Conelly, 1998) and in the basement drilled immediately east of the FEAAC in the Baltic Sea (Tschernoster, 2000; Fig. 5). This suggests that Baltican crust extends below the overthrust FEAAC and SJY. Based on seismic data, Winchester et al. (2002b) suggested that the Baltican basement of the western HEE is overlain by an Avalonian, 3- to 10-km thick thrust unit at a depth of 5–15 km, a view supported by the wider record of SHRIMP-dated inherited zircons reported by Breitkreuz et al. (2004). The HEE is separated from the East European craton by a large fault zone that also cuts through the overlying FEAAC and SJY thrust units (Fig. 5; cf. Dallmeyer et al., 1999). This fault is probably a Caledonian strike-slip fault, but the displacement may not be large.
North Sea Basement Terrane (NSB)
FAR EAST AVALONIAN TERRANES
The metamorphic rocks of the RZ south of Kraków in southern Poland (RZ in Fig. 1), entrained between the southern USL and the MPK, may represent a displaced fragment of Far East Avalonian crust. Conventional zircon dating on abraded zircons yielded Pb-Pb ages of 2.51–2.67 Ga for amphibolites and best fit–concordia intercepts at 2732 +23/–21 Ma and 1170 +59/– 60 Ma for a leucocratic vein; K-Ar dating yielded a cluster at 460–430 Ma for metamorphic overprinting (Bylina et al., 2000, and references therein). These results point to Archean rocks with a possible Grenvillian overprint, otherwise unknown on the southwestern Baltican margin (see earlier discussion). Metamorphic overprinting during the Late Ordovician or the Early Silurian has
This group summarizes terranes of Avalonian affinities that docked to the southwestern margin of Baltica during Late Ordovician time and are now detached from East Avalonia (Fig. 5; cf. Winchester et al., 2002a). Far East Avalonian Accretionary Complex (FEAAC) The FEAAC forms a thrust sheet on Baltican crust (Fig. 5; Winchester et al., 2002b). It was drilled in a 3130-m thick intersection on Rügen Island and consists of Tremadocian to Middle
The drilled basement of the North Sea, except for the Baltican basement area immediately west of Denmark (Fig. 5), records medium- to high-grade metamorphic cooling ages of ca. 450–435 Ma (K-Ar; Frost et al., 1981; Katzung, 2001), suggesting Late Ordovician crustal thickening during the docking of Far Eastern Avalonia to Baltica (Fig. 6). Similar low-grade metamorphic K-Ar and Ar-Ar ages have been recorded from the FEAAC (Dallmeyer et al., 1999; Katzung, 2001). With reference to the British Isles and the Brabant Massif (Fig. 1), a remnant oceanic zone in the current North Sea area is inferred by Verniers et al. (2002) to explain East Avalonian subduction-related volcanism from Ashgillian to Early Wenlockian time (Fig. 6). This volcanism is conventionally U-Pb zircon-dated to 433 ± 10 and 414 ± 16 Ma in the Brabant Massif (sills or necks; André and Deutsch, 1984) and to 426 +14/–15, 433 +9/–7, and 442 ± 22 Ma in the South Rhenish Massif (Sommermann et al., 1992, 1994). Accordingly, Figure 7 shows an oceanic area that is subducted under East Avalonia and replaced by the strike-slip invading NSB, previously deformed and metamorphosed at ca. 450 Ma during accretion to Southwest Baltica. Rzeszotary Terrane (RZ)
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not been recorded in the now neighboring USL and MPK, suggesting that the RZ does not form the basement to any of those. Istanbul-Zonguldak Terrane (IZ) The IZ contains Neoproterozoic metasediments, supra–subduction zone ophiolites, and meta-arc tonalites and granites with magmatic Pb-Pb zircon ages of 590–560 Ma (conventional and evaporation studies; Chen et al., 2002). Metasandstone from this basement records Rb-Sr ages of 548 ± 5 Ma (biotite–whole rock), 547 ± 5 (biotite-apatite) Ma, and 545 ± 5 Ma (biotite–apatite–whole rock; Chen et al., 2002). Nine detrital zircon grains from a metaquartzite show evaporation ages of 1862 ± 7 Ma, 1526 ± 3 Ma, 1436 ± 18 Ma, 1291 ± 6 Ma, 1204 ± 5 Ma, 1123 ± 19 Ma, 1166 ± 5 Ma, 884 ± 10 Ma, and 864 ± 4 Ma (Chen et al., 2002; only subtle overgrowths due to quartzitic sample composition; Pb hosted in potential overgrowths was removed by stepwise evaporation until the Pb-Pb ratios remained constant). This age spectrum is compatible with an Amazonian cratonic source (cf. Murphy et al., 2004). The 884–864 Ma zircons could be derived from rocks related to the Brasília fold belt, which contains arcs formed during 900–850 Ma, eventually accreted to eastern Amazonia at ca. 600 Ma (Pimentel et al., 2000; Da Silva et al., 2004). Zircon-dated magmatic rocks of ca. 850 Ma are also known from the northern margin of South America (Southwest Colombia; Aleman and Ramos, 2000). A relation to the West Baltican 800–880 Ma metamorphic zones of Denmark (Fig. 5) and South Norway (cf. Mulch and Cosca, 2004) is unlikely, because these zones record largely low-grade metamorphic overprinting, but no igneous activity of that age. In the eastern IZ, basement rocks are overstepped by over 800 m of Tremadocian–Arenigian shales and sandstones (Dean et al., 2000), similar to the CD and SD of East Moesia. A Late Arenigian trilobite assemblage is most comparable to such assemblages in East Avalonian Wales and again indicates location of the IZ in a peri–North Gondwana environment (Dean et al., 2000). Caradocian bedded carbonates, atypical for the North Gondwanan record, indicate location in a temperate climatic zone, as recorded on Southwest Baltica (cf. Modlinski et al., 1994), and the associated Caradocian trilobite assemblage is again most comparable to assemblages of East Avalonia (Dean et al., 2000). The western IZ (Istanbul region) records >1000 m of conglomerates, arkosic sandstones with fresh feldspars, and mudstones with reddish colors grading into overlying, fossildated Lowermost Silurian rocks (Fig. 4; Görür et al., 1997). This sequence is very probably of Upper Ordovician age and records erosion of a nearby uplifted basement area following docking of the IZ to the temperate shelves of Southwest Baltica. The entire IZ succession was variably deformed during Variscan collision, but the details are difficult to assess for superimposed Alpine deformation in the region (Görür et al., 1997). Variscan deformation suggests a more external position of the IZ on the South Laurussian margin (Figs. 6–9), similar to the situation found on the southern WM. Before the Middle
Cretaceous opening of the Black Sea basin (Nikishin et al., 2001) and the separation of the IZ from the Dobrogean terranes, the SD, CD, and IZ probably formed a coherent area (Fig. 9; Winchester et al., 2004). EASTERN AVALONIA Moravian Terrane (MV) The MV (Unrug et al., 1999) likely forms the continuation of East Avalonia (Fig. 5; cf. Friedl et al., 2000), wrapped around the Gondwana-derived part of the Bohemian Massif during Carboniferous indentation. In the south, the eastern boundary of the MV lies east of the ca. 580 Ma Brno Granitoid. While Devonian sediments overstep the Brno Granitoid, deep drilling immediately to the east revealed the presence of >1200-m thick, undeformed Lower Cambrian sandstones and shales (Vavrdova, 2004) typical of the Upper Silesia terrane. Inherited zircons in Moravian ca. 580 Ma granitoids yield SHRIMP-dated ages of 2.5 Ga, 2.0 Ga, 1.8–1.65 Ga, 1.5 Ga, and 1.2 Ga, favorably comparing to the inherited age spectrum of West Avalonian terranes and to the crustal components of Amazonia (Friedl et al., 2000). In the MV, the Upper Neoproterozoic basement is overstepped by Praguian conglomerates and sandstones that develop into a thick Middle–Upper Devonian arc-volcano-sedimentary succession (Patocˇka and Valenta, 1996), supporting a rifted arc environment for the Rhenohercynian basin on the South Laurussian margin (Fig. 8; Oczlon, 2006). Carpathian Lower Paleozoic Terrane (CLP) The East and South Carpathians contain Lower Devonian– lower Carboniferous overstep sequences, in places mimicking the Rhenohercynian volcano-sedimentary development of Germany and Moravia (Kräutner, 1997, for Romania). The Gemerides of Slovakia (GEM in Fig. 5; Vozárová and Vozár, 1997) and the Serbo-Macedonian zone (Karamata et al., 1997) are included here. These successions, absent from the contemporaneous North Gondwana margin, indicate that the Carpathian lower Paleozoic rocks were already located on the South Laurussian margin by Early Devonian time, and the entire area is referred to as the Carpathian lower Paleozoic terrane (CLP). Baltican and Avalonian affinities for the South Carpathians are suggested by shelly fauna in Lower Silurian sandstones, overstepping metamorphic rocks (Iordan and Sta˘noiu, 1993). This is in accord with the development of common Avalonian-Baltican faunal traits, expected after Late Ordovician docking of Avalonia to Baltica. Key faunal evidence from Cambrian–Lower Ordovician sediments is crucially lacking, and statistically significant detrital or inherited zircon studies currently do not allow for a clear assignment of the Carpathian Proterozoic basement and lower Paleozoic rocks to Avalonia or Baltica. However, Lower Ordovician arc-volcano-sedimentary successions in the East and South Carpathians (the Tulgeş Group; Kräutner, 1997; Vozárová and
Avalonian and Baltican terranes in the Moesian Platform Vozár, 1997; Munteanu et al., 1999; Vaida, 1999; Iancu and Berza, 2004) are unknown on the southwestern Baltican margin and may correlate with the East Avalonian record (Munteanu and Tatu, 2003; cf. Woodcock, 2000). This assumption is favored for the present review (Fig. 5), although parts of the outlined CLP may well be of Baltican origin, which remains to be shown by future studies of Early Ordovician microfossils and spot-dated detrital zircons. The earliest geochronological record from the South Carpathians indicates formation of primitive arc tholeiites about 1.57 Ga (143Nd/144Nd against 147Sm/144Nd isochron; Dra˘gus¸anu and Tanaka, 1999). Conventional U-Pb zircon dating on juvenile meta-arc-volcanic rocks in the same region yielded an emplacement age of 777 +/–3 Ma and Nd model ages of 817–717 Ma (Liégois et al., 1996). Intrusive ages of 588–567 Ma for K calcalkaline intrusives are reported for a different structural unit in that region (conventional U-Pb zircon), and Liégois et al. (1996) stress the overall similarity to the Saharan Pan-African record. Zincenco (1995) summarizes radiometric age data on the East and South Carpathians from papers published in Romanian and Russian journals and from unpublished Romanian state geological reports. No details other than formation names, dating methods, and the dated minerals are given. Although the quality of these data can hardly be assessed, they give a variety of clustered ages that comply with an East Avalonian framework. Three Rb-Sr whole-rock ages (681 ± 8 Ma, 677 ± 8 Ma, and 676 ± 4 Ma) and one K-Ar age on biotite (665 Ma), all from different areas, appear to record a metamorphic event that is also present in the basement of Central Dobrogea (696–643 Ma; see earlier discussion). Two zircon ages of 586 ± 7 Ma (U-Pb discordia) and 585 ± 15 Ma (PbPb) from the Novaci Granitoid (South Carpathians, NV in Fig. 5) may reflect the intrusive age. Rb-Sr whole-rock ages of 575 ± 3 Ma (paired with a K-Ar biotite age of 575 Ma), 575 ± 4 Ma, and 550 ± 26 Ma, all from metamorphic rocks in the South Carpathians, may date Latest Neoproterozoic metamorphism. Late Ordovician zircons in metamorphic rocks from the Tulghes¸ Group (East Carpathians) yielded dates of 454 ± 7 Ma (U-Pb discordia) and 451 ± 10 Ma (Pb-Pb discordia). Granitoids from the same group yielded zircon ages of 458 ± 7 Ma (U-Pb discordia) and 438 ± 7 Ma (Pb-Pb discordia). These ages form a distinctive cluster in the data set and may relate to Late Ordovician crustal thickening and melt production, which is supported by K-Ar ages of 472–450 Ma for metamorphic rocks of the Tulges¸ Group reported in Kräutner (1997, and references therein). Pana˘ et al. (2002) report conventional U-Pb concordia data on zircon from the Haghimas metagranitoid in the Bretila Unit (East Carpathians), with two lower intercept ages of 435.4 +/– 3.8 Ma and 433.3 +/– 3.3 Ma. A 450 Ma zircon age from a “stitching granitoid” in the Vlasina Complex in the Serbo-Macedonian zone of East Serbia could be genetically related (VL in Fig. 5; Karamata et al., 1997). Similar ages are recorded from the northern Far Eastern Avalonian terranes (NSB, FEAAC, and RZ, discussed earlier).
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In recognition of deformation and granitic melt generation by ca. 450 Ma, the CLP is shown in collision and crustal thickening during the docking of Far Eastern Avalonia with Southwest Baltica (Fig. 6). Another age group is formed by Rb-Sr wholerock ages of 420 ± 2 and 422 ± 9 Ma in metamorphic rocks (East Carpathians) and a K-Ar amphibole age of 420 Ma from the zircon-dated Upper Neoproterozoic Novaci Granitoid (585 Ma, South Carpathians). These Latest Silurian ages may be related to metamorphism in transpressive zones during the displacement of Caledonian terranes along the southwestern margin of Baltica. A strong Variscan overprint is evident from a cluster at 332–270 Ma recorded using all dating methods over the South and East Carpathians. More recent conventional dating of inherited zircons from the East Carpathians (Pana˘ et al., 2002) indicates the presence of U-Pb upper intercept ages of 1835 ± 8 Ma, 1917 ± 35 Ma, and 3048 ± 140 Ma. These zircons come from the basement of fossil-dated Lower Ordovician arc-volcano-sedimentary rocks (the Tulghes¸ Goup). Munteanu and Tatu (2003) interpret all of these units as Avalonia-related. Structurally above the Tulghes¸ Group appear metamorphic rocks with upper intercept zircon ages of 500–600 Ma and 1100–1200 Ma (the Haghimas metagranitoid of the Bretila Unit; Pana˘ et al., 2002), interpreted by Munteanu and Tatu (2003) as Southwest Baltica–derived. However, there is no evidence of a Grenvillian process along the southwestern Baltican margin (see earlier discussion), and the paired Grenvillian–Late Neoproterozoic ages correspond exactly to the record of some Avalonian terranes (cf. Murphy et al., 2000; Keppie et al., 2003). We propose that all of the East and South Carpathian units are related to the Cambrian–Early Ordovician North Gondwana margin as part of East Avalonia. The Lower Ordovician Tulghes¸ Group is best interpreted as the fill of a rifted arc on continental crust; the Bretila Group occupied the frontal, oceanward part of the rifted arc, emplaced on the intra-arc basin following Late Ordovician collision with Southwest Baltica. Attempting to restore the CLP to its likely pre-Carboniferous position (Fig. 8) suggests an external location on Southwest Baltica with respect to other Caledonian terranes, which is corroborated by Devonian volcano-sedimentary active margin formations and the intense Variscan deformation found in Moravia. STYLE AND TIMING OF CALEDONIAN TERRANE DISPLACEMENT Deformation recorded at about Late Ordovician time on the southern margin of the WM and on the LZW suggests that these terranes record the docking of Far East Avalonia to Southwest Baltica. The fact that minor ensuing deformation and uplift or none is recorded on many of the other Baltican terranes reviewed here suggests that docking and further displacement of Far East Avalonia along the southern part of the southwestern Baltican margin was highly oblique or margin-parallel. In contrast, the record in the north, with the development of a large thrust sheet (the FEAAC) and Late Ordovician metamorphic ages (the NSB
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and FEAAC) suggests a collisional scenario with crustal thickening (Fig. 6). This collision is also recorded in more distant parts of East Avalonia, where it produced mild deformation (the “Shelveian event”; Toghill, 1992). Some proximal Baltican terranes (ŁG, MPK, and LZW) record thick Upper Silurian, mostly Ludlovian, turbidite-shale successions, locally in excess of 1000 m (e.g., Belka et al., 2000; Malec, 2002). These sediments were probably deposited in a transtensional setting next to transpressionally uplifted and eroded areas. The apparently small extent of related basins, normal to strike, and the absence of any significant crustal thickening around these terranes does not support a flexural foreland setting for these deposits. At the margins of the MPK, Upper Silurian deposits were deformed in a transpressive setting before deposition of the Lower Devonian overstep sequence (Unrug et al., 1999), suggesting that transtension turned to transpression around the MPK during the Latest Silurian. An unconformity at the base of the Middle Devonian has been reported from the southern WM, with limestone conglomerates overstepping weakly eroded and inclined Lower Devonian sediments (Yanev and Boncheva, 1995, 1997). Associated uplift led to coarsening of the sedimentary input on the WM and SD at about the Early–Middle Devonian boundary. The same unconformity has been reported from the East IZ (Fig. 4; Derman, 1997). Both of these occurrences were located in an outboard position on Southwest Baltica with respect to the other regions discussed in this review (cf. Fig. 8), where this unconformity has not been reported. Therefore, it may relate to tectonic processes at the boundary of the southerly adjacent Rheic Ocean. On a continental scale, it may be related to the widespread Acadian deformation recorded along the southern margin of Laurussia (e.g., Keppie, 1993; Verniers et al., 2002). Post-Caledonian Overstep Sequence Deposits of a Middle–Upper Devonian carbonate platform overlie most Caledonian terranes (Fig. 4). The commonly 1000– 2500-m thick Middle Devonian–lower Carboniferous carbonates on the Moesian Platform are quite comparable to the equally thick and contemporaneous carbonate platform development in Moravia (eastern Czech Republic; cf. Vinogradov and Popescu, 1984; Hladil et al., 1999), the MPK, and the USL (Unrug et al., 1999). All of these areas were communicating on the same continental margin (Fig. 8), the thick Devonian carbonates sealing terrane boundaries active during Late Ordovician–Silurian and Early Devonian time. Toward the outer South Laurussian margin, the platform carbonates grade into volcano-sedimentary successions of a rifted arc environment (cf. Patocˇka and Valenta, 1996). VARISCAN DISPLACEMENT AND OVERPRINTING Variscan (Carboniferous–Early Permian) amalgamation of Laurussia, the Asian protocontinents, and the Gondwanaderived Hun superterrane (Stampfli, 2000) resulted in the
formation of Laurasia (the northern part of Pangea). Caledonian terranes that occupied an outer position on the Devonian South Laurussian margin, facing the Rheic Ocean to the south, record locally severe Carboniferous deformation, but metamorphism commonly does not exceed low-grade conditions. This is well documented in the MV, CLP, and southern WM, and to some extent in the IZ. The contact zone of the MPK and the ŁG (Holy Cross Mountains, South Poland) forms an exception, recording an intense early north-south to NNE-SSW compressional stage in an apparent hinterland location during Carboniferous deformation (Lamarche et al., 1999). Variscan indentation of the early Carboniferous Gondwana-derived part of the Bohemian Massif into the South Laurussian margin led to significant modifications. The Bohemian indenter is principally a collage of North Gondwanan proximal terranes, strike-slip juxtaposed during Late Devonian–early Carboniferous time in a dextral transcurrent regime of North Gondwana against the Rheic Ocean (Oczlon, 2006). The terranes so juxtaposed were widely thrust to the south and west over the northern and eastern margins of the Bohemian Massif (e.g., Zulauf et al., 2002) upon collision with South Laurussia. Ongoing convergence with South Laurussia led to indentation of the North Gondwanan crust thus amalgamated. The apparent wrapping of East Avalonia around the Bohemian indenter suggests Carboniferous rotation and/or escape tectonics (Fig. 5). Dextral Variscan transpression along the eastern margin of the Bohemian Massif, documented in Moravia by Schulmann and Gayer (2000), is opposed by sinistral transpression in the east along the concealed boundary between the USL and the MPK, affecting Middle–Upper Devonian carbonates (drillhole data; Fig. 10 in Unrug et al., 1999). This Variscan fault pattern reflects an escape structure, leading to the detached position and triangular shape of the Moravia–Upper Silesia region (USL). Within this scenario, ESE-directed escape is proposed for the MPK and the ŁG (Fig. 5). The dominant north-south to NNE-SSW compression recorded at the MPK-ŁG contact zone in the early stage of Carboniferous deformation (Lamarche et al., 1999) can be best explained if these terranes had a more westerly position in Early Carboniferous time, perhaps immediately north of the Bohemian indenter. Ensuing ESE-directed escape along the old Caledonian terrane boundaries brought the MPK and the ŁG into their current location, in an area where intense Variscan north-south-directed deformation would not be expected. The southern margin of the WM was deformed during Variscan (Carboniferous) accretion of the Gondwana-derived Balkan terrane to Laurussia (Fig. 9; Haydoutov and Yanev, 1997), and deformed Silurian–Devonian successions on the southern WM are recorded in several drillholes in North Bulgaria (Yanev and Boncheva, 1997). Repetition of conodont-dated Middle Devonian carbonates, with tectonic intercalation of Upper Devonian carbonates, is recorded in the Vaklino hole (VA in Fig. 2; Yanev and Boncheva, 1995). The deformed Paleozic rocks are in some drillholes overstepped by Permian clastic sediments (e.g., VE and DG in Fig. 2).
Avalonian and Baltican terranes in the Moesian Platform Mesozoic Displacement Because North Dobrogea was accreted to Baltica during Carboniferous time, later than Central Dobrogea, North Dobrogea must have been located more externally or at least laterally along Baltica with respect to Central Dobrogea during Carboniferous–Early Permian time (Fig. 9). The most likely scenario is thus a Mesozoic southeast-directed sinistral displacement along the East European craton, accommodated by the three Alpine faults that separate the Dobrogean terranes (Fig. 5). Mesozoic transpressional strike-slip displacement led to the juxtaposition of Central Dobrogea next to North Dobrogea along the Peceneaga-Camena Fault (Banks and Robinson, 1997), which is overstepped by Upper Cretaceous sediments (Fig. 3; Seghedi, 1998). Faulting affects Upper Jurassic (Oxfordian) carbonates. Perhaps simultaneously, the South Dobrogea-Palazu-West Moesian Platform block was juxtaposed next to Central Dobrogea along the 3-km wide Capidava-Ovidiu Fault Zone (cf. Figs. 3 and 10). However, this fault zone was also active during Late Cretaceous– Cainozoic convergence and collision (Figs. 2 and 3; Avram et al., 1998), when it assumed its current dextral structural fabric (cf. Hippolyte, 2002). Late Jurassic–Early Cretaceous reverse faulting and thrusting is observed around the Sfântu Gheorghe Fault (Fig. 3; North Dobrogea and the adjacent Pre-Dobrogean Depression; Belov et al., 1987), suggesting that the ND also records post-Variscan displacement against the East European craton. The ultimate cause for these displacements was the opening of basins floored by oceanic crust in the area of the present Carpathian orogen and around the Moesian Platform during Middle Jurassic–Early Cretaceous time (Fig. 10; Ceahla˘u-Severin Ocean, Csontos and Vörös, 2004; “Proto-Pannonian” oceanic basin, Banks and Robinson, 1997). Ophiolites of the Ceahla˘u-Severin Ocean are preserved in the East Carpathians (CE in Fig. 5) and in the South Carpathians, at the margin of the western Moesian Platform (SV in Fig. 5; Savu, 1985). There are no ophiolites farther to the south, but Upper Jurassic–Cretaceous outer shelf and slope sedimentation (turbidites, basinal marls) is preserved on the southern margin of the Moesian Platform and the northern margin of the Balkan terrane (cf. Tchoumatchenco and Sapunov, 1994; Minkovska et al., 2002). The Moesian Platform was detached from the area of the USL and the MPK during the opening of the Ceahla˘u-Severin Ocean and displaced ~500 km southeast along the East European craton, as suggested by the current northwest-southeast width of the Carpathian orogen. The Ceahla˘u ophiolites record obduction in late Early Cretaceous time, when crustal shortening and nappe stacking is recorded in the Carpathians and beyond. However, late Early Cretaceous deformation is due to the closure of an oceanic basin sutured in the current Inner Carpathian region (Csontos and Vörös, 2004), and Ceahla˘u-Severin oceanic crust was only marginally obducted, while the ocean remained open. This is implicit from Paleogene–early Neogene oceanic subduction under the Inner Carpathian region (Seghedi et al., 2004a) and underlined by lack of any crustal thickening around the Moesian Platform
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during late Early Cretaceous time (cf. Savu, 1985, and Minkovska et al., 2002). The Ceahla˘u-Severin Ocean was subducted during Late Cretaceous time under the Balkan terrane of northern Bulgaria (Fig. 10) and the Banat area in the South Carpathians (BA in Fig. 5). Upper Cretaceous arc-magmatic rocks in the Severin area are medium-K arc basalts and andesites with tholeiitelike geochemistry (SV in Fig. 5; Savu et al., 1987). Farther west, in the Banat area, Upper Cretaceous arc-magmatic rocks are of mostly calc-alkaline, high-K calc-alkaline to shoshonitic compositions (Ciobanu et al., 2002). This trend of increasing alkalinity from the Severin to the Banat area is consistent with west-directed subduction of the Ceahla˘u-Severin Ocean (present-day orientation; see Fig. 10). A steep subduction zone relatively near to the arc is suggested by the often superimposed magmatism of different alkalinity (cf. the compilation in Ciobanu et al., 2002) and by the narrowness of the Late Cretaceous intra-arc basin, which barely exceeds 100 km (deformation considered). These parameters contrast with models that derive the Upper Cretaceous arc-magmatic rocks from a remote, shallow-dipping, north-directed subduction zone in the south (e.g., Ciobanu et al., 2002; Csontos and Vörös, 2004). The northern part of the Ceahla˘u-Severin Ocean was subducted during Paleogene and early Neogene times under the current Inner Carpathian region (the remnant Carpathian oceanic basin of I. Seghedi et al., 2004a). CONCLUSIONS Several conclusions can be drawn from this chapter: • Four distinct units with a Caledonian displacement history can be outlined on the Moesian Platform; these are interpreted from west to east as the West Moesia (WM), South Dobrogea (SD), Palazu (PZ), and Central Dobrogea (CD) terranes. • The occurrence of Baltican early Middle Cambrian trilobites and a typically West Baltican succession of Cambrian facies suggest that the WM represents a proximal, Baltica-derived terrane broken off and displaced along the Southwest Baltican margin during the Late Ordovician– Silurian collision with Far East Avalonia. An Uppermost Neoproterozoic basement suggests provenance from the southern Baltican margin, probably next to the Scythian Platform (Figs. 1 and 6). • Based on the occurrence of Early Ordovician cold-water acritarchs, the SD was located at the North Gondwana margin during that time. Uppermost Neoproterozoic sediments in the CD were deposited in a fore-arc position, fed from a contemporaneous magmatic arc to the south with Rondonian, Grenvillian, and Upper Neoproterozoic sources unrelated to Baltica, but likely related to northern South America. Both SD and CD are interpreted as part of Far East Avalonia, accreted to Southwest Baltica during Late Ordovician time. • The 1.78–1.62 Ga basement encountered in drillholes in the PZ probably represents a sliver of Baltican crust.
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While old K-Ar data (Giusca et al., 1967; Mînzatu et al., 1975; re-calculated in Kräutner et al., 1988) need confirmation, rocks of that metamorphic age typically appear in West Baltica, but not in Avalonian regions. In addition, the overlying Neoproterozoic alkaline volcano-sedimentary Cocos¸u Formation is well correlatable with the West Baltican Neoproterozoic rift formations (cf. Bakun-Czubarow et al., 2002). Middle–Upper Silurian facies and shelly faunal associations in carbonates indicate a position on the Southwest Baltican margin for all Moesian terranes (Iordan, 1999). Unconformities at the base of and within the Silurian and at the base of the Devonian are characteristic of Caledonian displaced terranes along the southwestern margin of Baltica, linking the Moesian terranes with terranes located farther north, such as the Małopolska terrane (South Poland). The juxtaposition of terranes of Far East Avalonian and proximal Baltican origin is explained with postdocking, dextral strike-slip along the southwestern margin of Baltica during Late Ordovician–Silurian time (Figs. 6 and 7). In contrast, docking in the northern part of the southwestern Baltican margin (Northeast Germany and Poland) involved collision with metamorphic overprinting and thrusting of the Far East Avalonian accretionary complex over the southwestern Baltican margin. The Variscan indentation of the North Gondwana–derived Bohemian Massif into the South Laurussian margin and, more important, the Mesozoic breakup of southern Europe disrupted some Caledonian terranes (Fig. 10) and led to sinistral strike-slip displacements. During the Mesozoic opening of the Ceahla˘u-Severin Ocean in the current Carpathian region, the Moesian Platform was separated from the Upper Silesia and Małopolska terranes of South Poland. With respect to the East European craton, the estimated amount of Mesozoic sinistral southeast displacement of the Moesian Platform is 500 km, corresponding to the modern southeast-northwest extent of the injected Carpathian region.
ACKNOWLEDGMENTS We are thankful to Thomas Servais for advice on the distribution and assignment of Early Ordovician acritarchs and to Florentin Paris for commenting on the distribution of chitinozoans at about the Silurian–Devonian boundary. MO would like to thank Belevion Geological and Geophysical Consulting (Bucharest) for their search and supply of literature on the Moesian Platform. Constructive reviews significantly improving the quality of this paper were provided by John Winchester and Adrzej Z˙ elaz´niewicz. REFERENCES CITED Adamia, S.A., Belov, A.A., Chabukiani, A.O., Chkhotua, T.G., Lordkipanidze, M., and Shavishvili, I.D., 1997, Terrane description: The Caucasus: Annales Géologiques des Pays Helléniques, v. 37, p. 537–560.
Åhäll, K.-I., and Conelly, J., 1998, Intermittent 1.53–1.13 Ga magmatism in western Baltica: Age constraints and correlations within a postulated supercontinent: Precambrian Research, v. 92, p. 1–20, doi: 10.1016/ S0301-9268(98)00064-3. Åhäll, K.-I., Samuelsson, L., and Persson, P.-O., 1997, Geochronology and structural setting of the 1.38 Ga Torpa granite: Implications for charnockite formation in SW Sweden: Geologiska Föreningens i Stockholm Förhandlingar, v. 114, p. 449–450. Aleman, A., and Ramos, V.A., 2000, Northern Andes, in Cordani, U.G., et al., eds., Tectonic evolution of South America: 31st International Geological Congress, Rio de Janeiro, p. 453–480. André, L., and Deutsch, S., 1984, Les porphyres de Quenast et de Lessines: Géochronologie, géochimie isotopique et contribution au problème de l’âge du socle précambrien du Massif du Brabant (Belgique): Bulletin de la Société Belge de Géologie, v. 93, p. 375–384. Avram, E., Dumitru, I., and Dra˘goi, D., 1998, Capidava-Ovidiu fault system revisited: Revue Roumaine de Géologie, Géophysique et Géographie, Géologie, v. 42, p. 113–124. Bakun-Czubarow, N., Białowolska, A., and Fedoryshyn, Y., 2002, Neoproterozoic flood basalts of Zabolottya and Babino Beds of the volcanogenic Volhynian Series and Polesie Series dolerites in the western margin of the East European Craton: Acta Geologica Polonica, v. 52, p. 481–496. Banks, C.J., and Robinson, A.G., 1997, Mesozoic strike-slip back-arc basins of the Western Black Sea Region, in Robinson, A.G., ed., Regional and petroleum geology of the Black Sea Region: Tulsa, Oklahoma, American Association of Petroleum Geologists Memoir 68, p. 53–62. Barbu, C., Mehmet, N.A., and Paraschiv, C., 1970, Le Paléozo que de l’Avantpays des Carpates orientales entre la vallée de Buzau et la frontière septentrionale de la R.S.R: Acta Geologica Academiae Scientiarum Hungaricae, v. 14, p. 243–249. Beier, H., and Katzung, G., 2001, The deformation history of the Rügen Caledonides (NE Germany)—Implications from the structural inventory of the Rügen 5 borehole: Neues Jahrbuch für Geologie und Paläontologie, Abhandlungen, v. 222, p. 269–300. Beju, D., 1972, Zonare s¸i corelare a Paleozoicului din Platforma Moesica˘ pe baza asociaţiilor palino-protistologice: Petrol s¸i Gaze, v. 23, p. 7–15. Belka, Z., Ahrend, H., Franke, W., Schäfer, J., and Wemmer, K., 2000, The Baltica–Gondwana suture in central Europe: Evidence from K/Ar ages of detrital muscovites, in Franke, W., et al., eds., Orogenic processes: Quantification and modelling in the Variscan Belt: Geological Society, London, Special Publication 179, p. 87–102. Belka, Z., Valverde-Vaquero, P., Dörr, W., Ahrend, H., Wemmer, K., Franke, W., and Schäfer, J., 2002, Accretion of first Gondwana-derived terranes at the margin of Baltica, in Winchester, J.A., Pharaoh, T.C., and Verniers, J., eds., Palaeozoic amalgamation of Central Europe: Geological Society, London, Special Publication 201, p. 19–36. Belov, A.A., Slyusar, B.S., and Lartchenkov, E.P., 1987, Pre-Mesozoic formations of the north-west Black Sea region: Stratigraphic correlation forms and tectonics: International Geological Correlation Program no. 5: Newsletter 7, p. 119–128. Bingen, B., Nordgulen, Ø., Sigmond, E.M.O., Tucker, R., Mansfeld, J., and Högdal, K., 2003, Relations between 1.19–1.13 Ga continental magmatism, sedimentation and metamorphism, Sveconorwegian province: S Norway: Precambrian Research, v. 124, p. 215–241, doi: 10.1016/S03019268(03)00087-1. Bodzioch, A., Kozłowski, W., and Popławska, A., 2003, A Cooksonia-type flora from the Upper Silurian of the Holy Cross Mountains, Poland: Acta Palaeontologica Polonica, v. 48, p. 653–656. Bogdanova, S.V., Page, L.M., Skridlaite, G., and Taran, L.N., 2001, Proterozoic tectonothermal history in the western part of the East European Craton: 40 Ar/39Ar geochronological constraints: Tectonophysics, v. 339, p. 39–66, doi: 10.1016/S0040-1951(01)00033-6. Breitkreuz, C., and Kennedy, A., 1999, Magmatic flare-up at the Carboniferous/Permian boundary in the NE German Basin revealed by SHRIMP zircon ages: Tectonophysics, v. 302, p. 307–326, doi: 10.1016/S00401951(98)00293-5. Breitkreuz, C., Kennedy, A., Geissler, M., Ehling, B.-C., Kopp, J., Muszynski, A., Protas, A., and Stouge, S., 2004, Far East Avalonia: Its chronostratigraphic structure revealed by SHRIMP zircon data from Late Carboniferous volcanic rocks (drill cores from Germany, Poland and Denmark), in 32nd IGC, Florence, Scientific Sessions, Session 324, Abstracts (part 2), p.1432.
Avalonian and Baltican terranes in the Moesian Platform Brewer, T.S., Åhäll, K.-I., Menuge, J.F., Storey, C.D., and Parrish, R.R., 2004, Mesoproterozoic bimodal volcanism in SW Norway, evidence for recurring pre-Sveconorwegian continental margin tectonism: Precambrian Research, v. 134, p. 249–273, doi: 10.1016/j.precamres.2004.06.003. Brinkmann, R., and Krömmelbein, K., 1984, Abriss der Geologie, Historische Geologie: Stuttgart, Enke, 400 p. Buła, Z., and Jachowicz, M., 1998, Development of Lower Palaeozoic sediments in the foreland of the East European Platform (south-eastern Poland): Schriftenreihe des staatlichen Museums für Mineralogie und Geologie Dresden, v. 9, p.113–114. Bylina, P., Zelazniéwicz, A., and Dörr, W., 2000, Archean basement in the Upper Silesia Block: U-Pb zircon age from amphibolites of the Rzeszotary horst: Joint Meeting, EUROPROBE and PACE projects, Zakopane, Poland, Abstracts, p. 11–12. Cˇecˇys, A., Bogdanova, S., Janson, C., Bibikova, E., and Kornfält, K.-A., 2002, The Stenshuvud and Tåghusa granitoids: New representatives of Mesoproterozoic magmatism in southern Sweden: Geologiska Föreningens i Stockholm Förhandlingar, v. 124, p. 149–162. Chen, F., Siebel, W., Satir, M., Terzioglu, K., and Saka, K., 2002, Geochronology of the Karadere basement (NW Turkey) and implications for the geological evolution of the Istanbul Zone: International Journal of Earth Sciences, v. 91, p. 469–481, doi: 10.1007/s00531-001-0239-6. Ciobanu, C.L., Cook, N.J., and Stein, H., 2002, Regional setting and geochronology of the Late Cretaceous Banatitic Magmatic and Metallogenic Belt: Mineralium Deposita, v. 37, p. 541–567, doi: 10.1007/s00126-002-0272-9. Cocks, L.R.M., 2002, Key Lower Palaeozoic faunas from near the Trans-European Suture Zone, in Winchester, J.A., Pharaoh, T.C., and Verniers, J., eds., Palaeozoic amalgamation of Central Europe: Geological Society, London, Special Publication 201, p. 37–46. Cocks, L.R.M., and Fortey, R.A., 1998, The Lower Paleozoic margins of Baltica: Geologiska Föreningens i Stockholm Förhandlingar, v. 120, p. 173–179. Cocks, L.R.M., McKerrow, W.S., and Van Staal, C.R., 1997, The margins of Avalonia: Geological Magazine, v. 134, p. 627–636, doi: 10.1017/ S0016756897007425. Compston, W., Sambridge, M.S., Reinfrank, R.F., Moczydłowska, M., Vidal, G., and Claesson, S., 1995, Numerical ages of volcanic rocks and the earliest faunal zone within the Late Precambrian of Poland: Journal of the Geological Society, London, v. 152, p. 599–611. Crowley, Q.G., Marheine, D., Winchester, J.A., and Seghedi, A., 2000, Recent geochemical and geochronological studies in Dobrogea, Romania: TESZ and PACE Conference, Zakopane, Poland, Abstracts, p. 16–17. Csontos, L., and Vörös, A., 2004, Mesozoic plate tectonic reconstruction of the Carpathian region: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 210, p. 1–56, doi: 10.1016/j.palaeo.2004.02.033. Dallmeyer, R.D., Giese, U., Glasmacher, U., and Pickel, W., 1999, First 40Ar/39Ar age constraints for the Caledonian evolution of the Trans-European suture zone in NE Germany: Journal of the Geological Society, London, v. 156, p. 279–290. Da Silva, L.C., McNaughton, N.J., Armstrong, R., Hartmann, L.A., and Fletcher, I.R., 2005, The neoproterozoic Mantiqueira Province and its African connections: A zircon-based U–Pb geochronologic subdivision for the Brasiliano/Pan-African systems of orogens: Precambrian Research, v. 136, p. 203–240, doi: 10.1016/j.precamres.2004.10.004. Dean, W.T., Monod, O., Rickards, R.B., Demir, O., and Bultynck, P., 2000, Lower Palaeozoic stratigraphy and palaeontology, Karadere-Zirze area, Pontus Mountains, northern Turkey: Geological Magazine, v. 137, p. 555–582, doi: 10.1017/S0016756800004635. Derman, A.S., 1997, Sedimentary characteristics of Early Paleozoic rocks in the western Black Sea region: Turkish Association of Petroleum Geologists Special Publication 3, p. 25–38. Dra˘gus¸anu, C., and Tanaka, T., 1999, 1.57-Ga magmatism in the South Carpathians: Implications for the Pre-Alpine basement and the evolution of the mantle under the European continent: Journal of Geology, v. 107, p. 237–248. Dudek, A., and Melková, J., 1975, Radiometric age determination in the crystalline basement of the Carpathian foredeep and of the Moravian flysch: Bulletin of the Czech Geological Survey, v. 50, p. 257–264. Franke, D., 1994, The deformational history of the Caledonian terranes at Baltica’s southwestern margin: Zeitschrift für Geologische Wissenschaften, v. 22, p. 67–80. Friedl, G., Finger, F., McNaughton, N.J., and Fletcher, I.R., 2000, Deducing the ancestry of terranes: SHRIMP-evidence for South America–derived Gondwana fragments in Central Europe: Geology, v. 28, p. 1035–1038.
397
Frost, R.T.C., Fitch, F.J., and Miller, J.A., 1981, The age and nature of the crystalline basement of the North Sea Basin, in Illing, L.V., and Hobson, G.D., eds., Petroleum geology of the continental shelf of North-West Europe: London, Institute of Petroleum, p. 43–57. Gagała, Ł., 2005, Pre-Ordovician polyphase tectonics of the Cambrian sequences in the Kielce Unit, Holy Cross Mts. (Central Poland): Geological Quarterly, v. 49, p. 53–66. Garzanti, E., and Gaetani, M., 2002, Unroofing history of Late Palaeozoic magmatic arcs within the “Turan Plate” (Tuarkyr, Turkmenistan): Sedimentary Geology, v. 151, p. 67–87, doi: 10.1016/S0037-0738(01)00231-7. Giese, U., Handler, R., Neubauer, F., and Weber, J., 2001, 40Ar/39Ar-Laseranalyses of detrital white mica from the pre-Carboniferous of the Loissin 1 borehole—Implications for the deeper subsurface of NE-Germany: Neues Jahrbuch für Geologie und Paläontologie, Abhandlungen, v. 222, p. 215–251. Giusca, D., Ianovici, V., Manzatu, S., Soroiu, E., Lemne, M., Tanasescu, A., and Ioncica, M., 1967, Asupra varstei absolute a formatiunilor cristalofiliene din forlandul orogenului carpatic: Studii si Cercetari de Geologie, Geofizica˘, Geografie, seria Geologie, v. 12, p. 287–297. Glasmacher, U.A., Bauer, W., Giese, U., Reynolds, P., Kober, B., Puchkov, V.N., Stroink, L., Alekseyev, A.A., and Willner, A.P., 2001, The metamorphic complex of Beloretzk, SW Urals, Russia—A terrane with a polyphase Meso- to Neoproterozoic thermo-dynamic evolution: Precambrian Research, v. 110, p. 185–213, doi: 10.1016/S03019268(01)00187-5. Görür, N., Okay, A.I., Şengör, A.M.C., Tüysüz, O., Yiğitbas¸, E., Sakinç, M., Akkök, R., and Monod, O., 1997, Palaeogeographic and tectonic position of the Carboniferous rocks of the western Pontides (Turkey): Bulletin de la Société Géologique de France, v. 168, p. 197–205. Greiling, R.O., Samuelsson, L., and Weinert, H., 1988, Sedimentary dykes of Cambrian age in the Baltic Shield of SW Sweden (Lake Vänern to Gullmarn) and their tectonic implications, in Binzer, K., Marcussen, I., and Konradi, P., eds., 18. Nordiske Vintermode, Dansk Geologiske Undersøgelse, p. 143–144. Greiling, R.O., Jensen, S., and Smith, A.G., 1999, Vendian-Cambrian subsidence of the passive margin of western Baltica—Application of new stratigraphic data from the Scandinavian Caledonides: Norsk Geologisk Tidsskrift, v. 79, p. 133–144, doi: 10.1080/002919699433771. Gutiérrez-Marco, J.C., Yanev, S.N., Sachanski, V.V., Rabano, I., Lakova, I., San José Lanca, M.A., Díaz Martinez, E., Boncheva, I., and Sarmiento, N.G., 2003, New biostratigraphic data from the Ordovician of Bulgaria, in Albanesi, G.L., Beresi, M.S., and Peralta, S.H., eds., Ordovician from the Andes: INSUGEO, Serie Correlación Geologica, Tucumán, v. 17, p.79–85. Haydoutov, I., and Yanev, S., 1997, The Protomoesian microcontinent of the Balkan Peninsula—A peri-Gondwanaland piece: Tectonophysics, v. 272, p. 303–313, doi: 10.1016/S0040-1951(96)00264-8. Hippolyte, J.C., 2002, Geodynamics of Dobrogea (Romania): New constraints on the evolution of the Tornquist–Teisseyre Line, the Black Sea and the Carpathians: Tectonophysics, v. 357, p. 33–53, doi: 10.1016/S00401951(02)00361-X. Hladil, R., Melichar, R., Otava, J., Galle, A., Krs, M., Man, O., Pruner, P., Cˇejchan, P., and Orel, P., 1999, The Devonian in the easternmost Variscides, Moravia: A holistic analysis directed towards comprehension of the original context: Abhandlungen der geologischen Bundesanstalt: Wien, v. 54, p. 27–47. Huff, W.D., Bergström, S.M., and Kolata, D.R., 1992, Gigantic Ordovician volcanic ash fall in North America and Europe: Biological, tectonomagmatic and event stratigrafic significance: Geology, v. 20, p. 875–878, doi: 10.1130/0091-7613(1992)020<0875:GOVAFI>2.3.CO;2. Iancu, V., and Berza, T., 2004, Palaeozoic rock assemblages incorporated in the South Carpathians mobile belt overriding the Moesian Platform: Avalonia-Moesia Symposium and Workshop, Ghent/Ronse, Belgium, Abstracts, p. 19–20. Iordan, M., 1971, Graptoliţii din foragul Ţandarei (Platforma Morsica): Da˘ri de Seama˘ ale Şedinţelor, Institutul de Geologie s¸i Geofizica˘, v. 58, p. 31–47. Iordan, M., 1981, Study of Silurian and Devonian faunas from the Eastern part of the Moesian Platform: Memoriile Institutul de Geologie s¸i Geofizica˘, v. 30, p. 115–221. Iordan, M., 1984, Biostratigraphy of the Silurian and Devonian in the Moldavian and Moesian Platforms (Romania): Anuarul Institutului de Geologie s¸i Geofizica˘: Stratigrafie-Paleontologie, v. 64, p. 259–267.
398
Oczlon et al.
Iordan, M., 1988, Biostratigraphy of the Devonian in Romania, in Devonian of the world, v. 1: Proceedings of the 2nd International Symposium on the Devonian System, Calgary, Canada, 1987, p. 507–512. Iordan, M., 1990, A new contribution to the graptolite inventory from the Moesian Platform (Romania): Dari de Seama ale Şedinţelor, Institutul de Geologie s¸i Geofizica˘, v. 74, p.135–143. Iordan, M., 1999, Biostratigraphy of the Paleozoic in the Moesian Platform, excursion guide: Romanian Journal of Tectonics and Regional Geology, v. 77, suppl. no. 2, p. 18–23. Iordan, M., and Spasov, H., 1989, Lower Paleozoic (Cambrian–Silurian) sediments of the Moesian Platform: the XIVth Congress of the CarpathianBalkan Geological Association, Sofia, Abstracts, p. 702–705. Iordan, M., and Sta˘noiu, I., 1993, On the Lower Silurian macrofauna (Llandoverian) from the Valea Izvorului Formation (South Carpathians): Revue Roumaine de Géologie, Géophysique et Géographie, Géologie, v. 37, p. 63–76. Iordan, M., Iliescu, V., Visarion, A., Baltres, A., Sa˘ndulescu, E., and Seifert, K., 1985, Lito- and biostratigraphy of the Palaeozoic sequences in the Opris¸or and Gârla Mare boreholes (southwestern part of the Moesian Platform): Da˘ri de seama˘ ale Şedinţelor: Stratigrafie, v. 69, no. 4, p. 5–28. Kalvoda, J., Leichmann, J., Bábek, O., and Melichar, R., 2003, Brunovistulian terrane, Central Europe) and Istanbul zone, NW Turkey, Late Proterozoic and Paleozoic tectonostratigraphic development and paleogeography: Geologica Carpathica, v. 54, no. 3, p. 139–152. Kämpf, H., Korich, D., and Brause, H., 1994, Anorthosit-, Gabbro- und Ilmeniterz-Xenolithe in permischen Basalten der Tiefbohrung Schwerin 1—Ein petrographisches Indiz für das Ostelbische Massif: Zeitschrift für Geologische Wissenschaften, v. 22, p. 115–228. Karamata, S., Krstic, B., Dimitrijevic, D.M., Dimitrijevic, M.N., Kneev, V., Stojanov, R., and Filipovic, I., 1997, Terranes between the Moesian Plate and the Adriatic Sea: Annales Géologiques des Pays Helléniques, v. 37, p. 429–477. Katzung, G., 2001, The Caledonides at the southern margin of the East European Craton: Neues Jahrbuch für Geologie und Paläontologie, Abhandlungen, v. 222, p. 3–53. Kent, D.V., and Smethurst, M.A., 1998, Shallow bias of paleomagnetic inclinations in the Paleozoic and Precambrian: Earth and Planetary Science Letters, v. 160, p. 391–402, doi: 10.1016/S0012-821X(98)00099-5. Keppie, J.D., 1989, Northern Appalachian terranes and their accretionary history: Boulder, Colorado, Geological Society of America Special Paper 230, p. 159–192. Keppie, J.D., 1993, Synthesis of Palaeozoic deformational events and terrane accretion in the Canadian Appalachians: Geologische Rundschau, v. 82, p. 381–431, doi: 10.1007/BF00212406. Keppie, J.D., Nance, R.D., Murphy, J.B., and Dostal, J., 2003, Tethyan, Mediterranean, and Pacific analogues for the Neoproterozoic–Palaeozoic birth and development of peri-Gondwanan terranes and their transfer to Laurentia and Laurussia: Tectonophysics, v. 365, p. 195–219, doi: 10.1016/ S0040-1951(03)00037-4. Khain, V.E., 1994, Geology of Northern Eurasia (Ex-USSR): Beiträge zur regionalen Geologie der Erde, v. 24, 404 p. Khizhnjakov, A.V., 1969, The age of the old orogens surrounding the SW border of the Russian plate and their relationship to the platform: Geologija i geochimija gorjuecich iskopaemych, 20, p. 3–11 (in Russian). Kräutner, H.G., 1997, Alpine and Pre-Alpine Terranes in the Romanian Carpathians and Apuseni Mts: Annales Géologiques des Pays Helléniques, v. 37, p. 331–400. Kräutner, H.G., Muresan, M., and Seghedi, A., 1988, Precambrian of Dobrogea, in Zoubek, V., ed., Precambrian in younger fold belts: New York, John Wiley, p. 361–379. Lakova, I., 1995, Paleogeographical affinities of Pridolian and Lochkovian chitinozoans from North Bulgaria: Geologica Balcanica, v. 25, p. 23–28. Lamarche, J., Mansky, J.L., Bergerat, F., Averbuch, O., Hakenberg, M., Lewandowski, M., Stupnicka, E., Swidrowska, J., Wajsprych, B., and Wieczorek, J., 1999, Variscan tectonics in the Holy Cross Mountains (Poland) and the role of structural inheritance during Alpine tectonics: Tectonophysics, v. 313, p. 171–186, doi: 10.1016/S0040-1951(99)00195-X. Liégois, J.P., Berza, T., Tatu, M., and Duchesne, J.C., 1996, The Neoproterozoic Pan-African basement from the Alpine Lower Danubian nappe system (South Carpathians, Romania): Precambrian Research, v. 80, p. 281–301. Malec, J., 2002, Sedimentology of deposits from around the Late Caledonian unconformity in the western Holy Cross Mts: Geological Quarterly, v. 45, p. 397–416.
Malinowski, M., Z˙ elaz´niewicz, A., Grad, M., Guterch, A., Janik, T., and CELEBRATION Working Group, 2005, Seismic and geological structure of the crust in the transition from Baltica to Palaeozoic Europe in SE Poland—CELEBRATION 2000 experiment, profile CEL02: Tectonophysics, v. 401, p. 55–77. Marheine, D., and Valverde-Vaquero, P., 2002, Recognition of pre-Sveconorwegian cooling ages in the Eastern European Craton, Central Poland: New 40Ar/39Ar dating in the 1.8 Ga Kampinos Complex: Precambrian Research, v. 118, p. 169–177, doi: 10.1016/S0301-9268(02)00095-5. Masiak, M., Podhalan´ska, T., and Stempin´-Sałek, M., 2003, Ordovician–Silurian boundary in the Bardo Syncline, Holy Cross Mountains, Poland— New data on fossil assemblages and sedimentary succession: Geological Quarterly, v. 47, p. 311–330. Minkovska, V., Peybernès, B., Nikolov, T., and Ivanov, M., 2002, Paleogeographic reconstruction of a segment of the North-Tethyan margin in Bulgaria from Barremian to Albian: Eclogae Geologicae Helvetiae, v. 95, p. 183–195. Mînzatu, S., Lemne, M., Vîjdea, E., Ta˘na˘sescu, A., Ioncica˘, M., and Tiepac, I., 1975, Date geochronologice obţinute pentru formaţiuni cristalofiliene s¸i masive eruptive din România: Da˘ri de seama˘ ale Şedinţelor, Institutul de Geologie s¸i Geofizica˘, v. 56/5, p. 85–111. Mocanu, V., Radulescu, F., and Dinu, C., 1991, Lithospheric structure of Romania and recent crustal movements, geodynamic connections: Revue Roumaine de Géologie, Géophysique et Géographie, v. 35, p. 1–11. Moczydłowska, M., 1995, Neoproterozoic and Cambrian successions deposited on the East European Platform and Cadomian basement area in Poland: Studia Geophysica et Geodetica, v. 39, p. 276–285, doi: 10.1007/ BF02295819. Moczydłowska, M., 1997, Proterozoic and Cambrian successions in Upper Silesia: An Avalonian terrane in southern Poland: Geological Magazine, v. 134, p. 679–689, doi: 10.1017/S0016756897007504. Modlinski, Z., Nehring-Lehfeld, M., and Ryba, J., 1994, The Early Palaeozoic Complex in the Polish part of the Baltic Sea: Zeitschrift für Geologische Wissenschaften, v. 22, p. 227–234. Morgan, J.W., Stein, H.J., Hannah, J.L., Markey, R.J., and Wiszniewska, J., 2000, Re-Os studies of Fe-Ti-V oxide and Fe-Cu-Ni sulfide deposits, Suwalki Anorthosite Massif, northeast Poland: Mineralium Deposita, v. 35, p. 391–401, doi: 10.1007/s001260050251. Mulch, A., and Cosca, M.A., 2004, Recrystallization or cooling ages: In situ UV-Laser 40Ar/39Ar geochronology of muscovite in mylonitic rocks: Journal of the Geological Society, London, v. 161, p. 573–582. Munteanu, A., Rosu, E., and Voda, A., 1999, Caledonian metvolcanics from the Eastern Carpathians: Geochemistry and tectonic setting: Romanian Journal of Tectonics and Regional Geology, v. 77, p. 70–71. Munteanu, A., and Tatu, M., 2003, The East-Carpathian Crystalline-Mesozoic Zone (Romania): Paleozoic amalgamation of Gondwana- and East European Craton–derived terranes: Gondwana Research, v. 6, p. 185–196, doi: 10.1016/S1342-937X(05)70969-2. Murgeanu, G., and Patrulius, D., 1963, Le Paléozoïque de la Plateforme Moesienne: Bulletin du VIe Congrès de l’Association Géologique CarpathoBalkanique, v. 1, p. 237–244. Murphy, J.B., Strachan, R.A., Nance, R.D., Parker, K.D., and Fowler, M.B., 2000, Proto-Avalonia: A 1.2–1.0 Ga tectonothermal event and constraints for the evolution of Rodinia: Geology, v. 28, p. 1071–1074, doi: 10.1130/0091-7613(2000)028<1071:PAAGTE>2.3.CO;2. Murphy, J.B., Pisevarsky, S.A., Nance, R.D., and Keppie, J.D., 2004, Neoproterozoic–Early Paleozoic evolution of peri-Gondwanan terranes: Implications for Laurentia-Gondwana connections: International Journal of Earth Sciences, v. 93, p. 659–682, doi: 10.1007/s00531-004-0412-9. Muţiu, R., 1991, On the presence of the Cambrian in the Moesian Platform (Romania): Revue Roumaine de Géologie, Géophysique et Géographie, Géologie, v. 35, p. 75–85. Narkiewicz, M., 2002, Ordovician through earliest Devonian development of the Holy Cross Mts. (Poland): Constraints from subsidence analysis and thermal maturity data: Geological Quarterly, v. 46, p. 255–266. Natal’in, B.A., and Şengör, A.M.C., 2005, Late Palaeozoic to Triassic evolution of the Turan and Scythian platforms: The pre-history of the PalaeoTethyan closure: Tectonphysics, v. 404, p. 175–202, doi: 10.1016/j.tecto. 2005.04.011. ˙ ylin´ska, A., Buła, Z., Grabowski, J., Krzymiec, P., and Poprawa, Nawrocki, J., Z P., 2004, Early Cambrian location and affinities of the Brunovistulian terrane (Central Europe) in the light of palaeomagnetic data: Journal of the Geological Society, London, v. 161, p. 513–522.
Avalonian and Baltican terranes in the Moesian Platform Nikishin, A.M., Ziegler, P.A., Panov, D.I., Nazarevich, B.P., Brunet, M.-F., Stephenson, R.A., Bolotov, S.N., Korotaev, M.V., and Tikhmirov, P.L., 2001, Mesozoic and Cainozoic evolution of the Scythian platform–Black Sea–Caucasus domain: Peri-Tethys Memoir, v. 6: Mémoires du Musée Nationale de l’Histoire Naturelle, v. 186, p. 295–346. Oaie, G., 1998, Sedimentological significance of mudstone microclast intervals in Upper Proterozoic turbidites, central Dobrogea, Romania: Sedimentary Geology, v. 115, p. 289–300, doi: 10.1016/S0037-0738(97)00097-3. Oaie, G., Seghedi, A., and Vaida, M., 2004, Sedimentology and source areas for the Neoproterozoic–Lower Cambrian turbidites from East Moesia, in Avalonia-Moesia Symposium and Workshop, Ghent/Ronse, Belgium, Abstracts, p. 23–24. Obst, K., Hammer, J., Katzung, G., and Korich, D., 2004, The Mesoproterozoic basement in the southern Baltic Sea: Insight from the G14–1 off-shore borehole: International Journal of Earth Sciences, v. 93, p. 1–12, doi: 10.1007/s00531-003-0371-6. Oczlon, M.S., 2006, Terrane map of Europe (1st edition): Gaea Heidelbergensis, v. 15. Available only online at http://www.geobib.uni-hd.de/terranemap/ index.html. Olaru, L., 1999, Palynology of the formation with paradoxissimus from Mitrofani of the Moesian Platform, Romania: Analele Ştiinţifice ale Universita˘ţii “Al. I. Cuza”: Ias¸i, Geologie, v. 55–56, p. 150–160. Pamic´, J., Tomljenovic´, B., and Balen, D., 2002, Geodynamic and petrogenetic evolution of Alpine ophiolites from the central and NW Dinarides: An overview: Lithos, v. 65, p. 113–142, doi: 10.1016/S00244937(02)00162-7. Pana˘, D.I., Balintoni, I., Heaman, L., and Creaser, R., 2002, The U-Pb and Sm-Nd dating of the main lithotectonic assemblages of the East Carpathians, Romania: Geologica Carpathica, Bratislava, v. 53, Special Issue, p.177–180. Paraschiv, D., 1974, Studiul stratigrafic al Devonianului s¸i Carboniferului din Platforma Moesica˘, la vest de raul Arges: Studii Tehnice s¸i Economice: Institutul Geologic, v. 12, p. 1–165. Paraschiv, D., 1986, Privire de ansamblu asupra magmatitelor din Platforma Moesica˘: Mine, Petrol s¸i Gaze, v. 37, p. 344–350. Paraschiv, D., Demetrescu, C., and Soroiu, M., 1982, Date geocronologice referitoare la fundamentul metamorfic s¸i unele corpuri magmatice din platforma Moesica˘: Mine, Petrol s¸i Gaze, v. 33/5, p. 239–243. Paraschiv, D., Demetrescu, C., and Soroiu, M., 1983a, Noi date geocronologice privind fundamentul metamorfic s¸i la unele corpuri magmatice din platforma Moesica: Mine, Petrol s¸i Gaze, v. 34/2, p. 94–96. Paraschiv, D., Danet, N., Popescu, M., and Dumitrescu, V., 1983b, The present stage of pre-Jurassic deposits knowledge in South Dobrogea: Anuarul Institutului de Geologie s¸i Geofizica˘, v. 59, p. 29–37. Paris, F., and Grahn, Y., 1996, Chitinozoa of the Silurian–Devonian boundary sections in Podolia, Ukraine: Palaeontology, v. 39, no. 3, p. 629–649. Paris, F., and Nolvak, J., 1999, Biological interpretation and paleobiodiversity of a cryptic fossil group: The “chitinozoan animal”: Géobios, v. 32, p. 315–324, doi: 10.1016/S0016-6995(99)80045-X. Patocˇka, F., and Valenta, J., 1996, Geochemistry of the Late Devonian intermediate to acid metavolcanic rocks from the southern part of the Vrbno Group, the Jeseníky Mts. (Moravo-Silesian Belt, Bohemian Massif, Czech Republic): Palaeotectonic implications: Geolines, v. 4, p. 42–54. Pharaoh, T.C., 1999, Palaeozoic terranes and their lithospheric boundaries within the Trans-European Suture Zone, TESZ, a review: Tectonophysics, v. 314, p. 17–41, doi: 10.1016/S0040-1951(99)00235-8. Pimentel, M.M., Fuck, R.A., Jost, H., Ferreira Fihlo, F.F., and Araujo, S.M., 2000, The basement of the Brasília Fold Belt and the Golás Magmatic Arc, in Cordani, U.G., et al., eds., Tectonic evolution of South America, 31st International Geological Congress, Rio de Janeiro, p. 195–229. Ra˘ileanu, G., Iordan, M., and Sa˘ndulescu, E., 1967, Consideraţii aspra Paleozoicului inferior din zona Ca˘la˘ras¸i: Da˘ri de seama˘ ale Şedinţelor, Institutul de Geologie s¸i Geofizica˘, v. 53/1, p. 419–426. Rickards, R.B., and Iordan, M., 1975, Romanian graptolites from boreholes on the Moesian Platform: Geological Magazine, v. 112, p. 241–255 Robardet, M., 2003, The Armorica ‘microplate’: Fact or fiction? Critical review of the concept and contradictory palaeobiogeographical data: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 195, p. 125–148, doi: 10.1016/S0031-0182(03)00305-5. Roberts, D., and Siedlecka, A., 2002, Timanian deformation along the northeastern margin of Baltica, Northwest Russia and Northeast Norway, and Avalonian–Cadomian connections: Tectonophysics, v. 352, p. 169–184, doi: 10.1016/S0040-1951(02)00195-6.
399
Rushton, A.W.A., and McKerrow, W.S., 2000, Cambrian fossils from the Moesian Platform and their relationships to the Trans-European Suture Zone: TESZ & PACE Conference, Zakopane, Poland, Abstracts, p. 71–72. Samuelsson, J., Vecoli, M., Bednarczyk, W.S., and Verniers, J., 2002, Timing of the Avalonia–Baltica plate convergence as inferred from palaeogeographic and stratigraphic data of chitinozoan assemblages in west Pomerania, northern Poland, in Winchester, J.A., Pharaoh, T.C., and Verniers, J., eds., Palaeozoic amalgamation of Central Europe: Geological Society, London, Special Publication 201, p. 95–113. Savu, H., 1985, Tectonic position and origin of the Alpine ophiolites in the Mehedinţi Plateau (South Carpathians) with special regard to those from the Podeni-Isverna-Nadanova Region: Da˘ri de seama˘ ale Şedinţelor, Institutul de Geologie s¸i Geofizica˘, v. 64/5, p. 57–71. Savu, H., and Paraschiv, D., 1982, Contributions to the study of the pre-Triassic magmatites in the Moesian platform: Dari de seama ale Şedinţelor, Insitutul de Geologia s¸i Geofizica, v. 64, no. 5, p. 73–81. Savu, H., and Stoian, M., 1998, The Gâlmele Îns¸irate ocean crust slice—A vestige of the Paleozoic ophiolite suture in North Dobrogea: Revue Roumaine de Géologie, Géophysique et Géographie, Géologie, v. 42, p. 29–42. Savu, H., Udrescu, C., Lemne, M., Romanescu, O., Stoian, M., and Neacs¸u, V., 1987, Island arc volcanics related to the wildflysch on the outer margin of the Danubia autochthon (Southern Carpathians) and their geotectonic implications: Revue Roumaine de Géologie, Géophysique et Géographie, Géologie, v. 31, p. 19–27. Schulmann, K., and Gayer, R., 2000, A model for a continental accretionary wedge developed by oblique collision: The NE Bohemian Massif: Journal of the Geological Society, London, v. 157, p. 401–416. Seghedi, A., 1998, The Romanian Carpathian Foreland: Politechnika Warszawska: Reports on Geodesy, v. 7, no. 37, Monograph on the Southern Carpathians, p. 21–48. Seghedi, A., and Oaie, G., 1994, Tectonic setting of two contrasting types of pre-Alpine basement: North versus Central Dobrogea: Romanian Journal of Tectonics and Regional Geology, v. 75, p. 56. Seghedi, A., and Oaie, G., 1995, Palaeozoic evolution of North Dobrodgea, in IGCP Project No. 369, Comparative evolution of Peri-Tethyan Rift Basins, field guidebook, Central and North Dobrodgea: Bucharest, Geological Institute of Romania, p. 5–75. Seghedi, A., and Oaie, G., 1999, Overview of the geology of Dobrogea: Romanian Journal of Tectonics and Regional Geology, v. 77, suppl. 2, p. 1–17. Seghedi, A., Lang, B., and Heimann, A., 1999, The deformational history of North Dobrogean Hercynian basement as reflected in new 39Ar/ 40Ar determinations: Romanian Journal of Tectonics and Regional Geology, v. 77, suppl. 1, p. 64–65. Seghedi, A., Kasper, H.U., and Maruntiu, M., 2000, Neoproterozoic intraplate magmatism in Moesia: Petrologic and geochemical data: TESZ & PACE Conference, Zakopane, Poland, Abstracts, p. 76. Seghedi, A., Spear, F., and Storm, L., 2003, The metamorphic evolution of the Hercynian basement of North Dobrogea: Constraints from petrological studies and monazite dating: The Fourth Stephan Mueller Conference of the EGS, Romania, Abstracts, p. 82. Seghedi, I., Downes, H., Szakács, A., Mason, P.R.D., Thirlwall, M.F., Rosu, E., Pécskay, Z., Márton, E., and Panaiotu, C., 2004a, Neogene magmatism and geodynamics in the Carpathian–Pannonian region: A synthesis: Lithos, v. 72, p. 117–146, doi: 10.1016/j.lithos.2003.08.006. Seghedi, A., Oaie, G., Vaida, M., Debacker, T.N., and Sintubin, M., 2004b, Palaeozoic formations in North Dobrodgea: Sedimentation, deformation and metamorphism: Avalonia-Moesia Symposium and Workshop, Ghent/ Ronse, Belgium, Abstracts, p. 31–32. Semenov, V.Y., Jankowski, J., Ernst, T., Józ´wiak, W., Pawliszyn, J., and Lewandowski, M., 1998, Electromagnetic soundings across the Holy Cross Mountains: Acta Geophysica Polonica, v. 46, p. 171–185. Servais, T., and Fatka, O., 1997, Recognition of the Trans-European Suture Zone (TESZ) by the palaeobiogeographical distribution pattern of Early to Middle Ordovician acritarchs: Geological Magazine, v. 134, p. 617– 625, doi: 10.1017/S001675689700753X. Servais, T., Li, J., Molyneux, S., and Raevskaya, E., 2003, Ordovician organicwalled microphytoplankton (acritarch) distribution: The global scenario: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 195, p. 149–172, doi: 10.1016/S0031-0182(03)00306-7. Shchipansky, A.A., and Bogdanova, S.V., 1996, The Sarmatian crustal segment: Precambrian correlation between the Voronezh Massif and the Ukrainian
400
Oczlon et al.
Shield across the Dniepr-Donets Aulagogen: Tectonophysics, v. 268, p. 109–125, doi: 10.1016/S0040-1951(96)00227-2. Sommermann, A.-E., Meisl, S., and Todt, W., 1992, Zirkonalter aus drei verschiedenen Metavulkaniten aus dem Südtaunus: Geologisches Jahrbuch Hessen, v. 120, p. 67–76. Sommermann, A.-E., Anderle, H.-J., and Todt, W., 1994, Das Alter des Quartzkeratophyrs der Krausaue bei Rüdesheim am Rhein (Bl.6013 Bingen, Rheinisches Schiefergebirge): Geologisches Jahrbuch Hessen, v. 122, p. 143–157. Spassov, C., Tenchov, J., and Janev, S., 1978, Die paläozoischen Ablagerungen in Bulgarien: Österreichische Akademie der Wissenschaften: Erdwissenschaftlicher Kommissionsband, v. 3, p. 279–296. Stampfli, G.M., 2000, Tethyan Oceans, in Winchester, J.A., and Piper, J.D.A., eds., Tectonics and magmatism in Turkey and the surrounding area: Geological Society, London, Special Publication 173, p. 1–23. Steemans, P., and Lakova, I., 2004, The Moesian Terrane during the Lochkovian—A new palaeogeographic and phytogeographic hypothesis based on miospore assemblages: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 208, p. 225–233, doi: 10.1016/j.palaeo.2004.03.006. Strachan, R.A., Nance, R.D., Dallmeyer, R.D., D’Lemos, R.S., Murphy, J.B., and Watt, G.R., 1996, Late Precambrian tectonothermal evolution of the Malverns Complex: Journal of the Geological Society, London, v. 153, p. 589–600. Tari, G., Dicea, O., Faulkerson, J., Georgiev, G., Popov, S., Stefanescu, M., and Weir, G., 1997, Cimmerian and Alpine stratigraphy and structural evolution of the Moesian Platform (Romania/Bulgaria): American Association of Petroleum Geologists Memoir 68, p. 63–90. Tchoumatchenco, P., and Sapunov, I., 1994, Intraplate tectonics in the Bulgarian part of the Moesian Platform during the Jurassic: Geologica Balcanica, v. 24, p. 3–12. Toghill, P., 1992, The Shelveian event, a Late Ordovician tectonic episode in Southern Britain (Eastern Avalonia): Proceedings of the Geologists’ Association. Geologists’ Association, v. 103, p. 31–35. Torsvik, T.H., and Van der Voo, R., 2002, Refining Gondwana and Pangea palaeogeography: Estimates of Phanerozoic non-dipole (octupole) fields: Geophysical Journal International, v. 151, p. 771–794, doi: 10.1046/ j.1365-246X.2002.01799.x. Tschernoster, R., 2000, Isotopengeochemische Untersuchungen am Detritus der Dänisch-Norddeutsch-Polnischen Kaledoniden und deren Vorland [Ph.D. thesis]: Aachen, Germany, Berichte aus der Geowissenschaft, Shaker Verlag, 128 p. Unrug, R., Czesław, H., and Chocyk-Jamin´ska, M., 1999, Easternmost Avalonian and Armorican–Cadomian Terranes of Central Europe and Caledonian–Variscan evolution of the polydeformed Kraków Mobile Belt: Geological constraints: Tectonophysics, v. 302, p. 133–157, doi: 10.1016/ S0040-1951(98)00278-9. Vaida, M., 1999, Dating and correlation based on palynological data of crystalline formations and associated syngenetic mineralizations from the southern crystalline massif of Eastern Carpathians and Eastern South Carpathians [Ph.D. thesis]: Ias¸i, Romania, Ias¸i University, 122 p. Vaida, M., Seghedi, A., and Verniers, J., 2005, Northern Gondwana affinity of the East Moesian Platform based on chitinozoans: Tectonophysics, v. 410, p. 379–387, doi: 10.1016/j.tecto.2004.12.039. Vavrdova, M., 2004, The Brunovistulicum: Assumptions and data: Zeitschrift der deutschen geologischen Gesellschaft, v. 155/1, p. 1–9. Verniers, J., Pharaoh, T., André, L., Debacker, T., De Vos, W., Everaerts, M., Herbosch, A., Samuelsson, J., Sintubin, M., and Vecoli, M., 2002, The Cambrian to mid Devonian history of Eastern Avalonia, east of the Midlands Microcraton: New data and a review, in Winchester, J.A., Pharaoh, T.C.,
and Verniers, J., eds., Palaeozoic amalgamation of Central Europe: Geological Society, London, Special Publication 201, p. 47–93. Vinogradov, C., and Popescu, M., 1984, Les faciès carbonatés du Paléozoïque de la plate-forme Moesienne (Roumanie): Anuarul Institutului de Geologie s¸i Geofizica˘, v. 64, p. 317–322. Vozárová, A., and Vozár, J., 1997, Terranes of the West Carpathian–North Pannonian Domain: Annales Géologiques des Pays Helléniques, v. 37, p. 245–269. Wagner, H., 1977, Rumänien: Bundesanstalt für Geowissenschaften und Rohstoffe: Hannover, Rohstoffwirtschaftliche Länderberichte, v. 13, 176 p. Winchester, J.A., 2003, Fragmentation of microcontinents accreting to Baltica: Did it happen and are the detached terranes traceable?: Journal of the Czech Geological Society, v. 48, no. 1–2, p. 133. Winchester, J.A., Pharaoh, T.C., and Verniers, J., 2002a, Palaeozoic amalgamation of Central Europe: An introduction and synthesis of new results from recent geological investigations, in Winchester, J.A., Pharaoh, T.C., and Verniers, J., eds., Palaeozoic amalgamation of Central Europe: Geological Society, London, Special Publication 201, p. 1–18. Winchester, J.A., and the PACE TMR Network Team (contract ERBFMRXCT97–0136), 2002b, Paleozoic amalgamation of Central Europe: New results from recent geological and geophysical investigations: Tectonophysics, v. 360, p. 5–21. Winchester, J.A., Yiğitbas¸, E., Bozkurt, E., and Pharaoh, T.C., 2004, Microcontinent-promontory interactions: Is there a link between Avalonia, Moesia and the Istanbul Block, NW Turkey? Avalonia-Moesia Symposium and Workshop, Ghent/Ronse, Belgium, Abstracts, p. 47. Woodcock, N.H., 1990, Sequence stratigraphy of the Palaeozoic Welsh Basin: Journal of the Geological Society, London, v. 147, p. 537–547. Woodcock, N.H., 2000, Ordovician volcanism and sedimentation on Eastern Avalonia, in Woodcock, N.H., and Strachan, R.A., eds., Geological history of Britain and Ireland: Oxford, Blackwell Science, p. 153–167. Yanev, S.N., 1992, Contribution to the elucidation of pre-Alpine evolution in Bulgaria (based on sedimentological data from marine Paleozoic): Geologica Balcanica, v. 22, p. 3–31. Yanev, S.N., and Boncheva, I., 1995, Contribution to the Paleozoic evolution of the recent Moesian platform: Geologica Balcanica, v. 25, p. 3–21. Yanev, S.N., and Boncheva, I., 1997, New data on the collision between peri-Gondwana, the Moesian Terrane and Dobrudja periphery of PaleoEurope: Turkish Association of Petroleum Geologists, Special Publication 3, p. 118–132. Z˙ elaz´niewicz, A., Seghedi, A., Jachowicz, M., Bobiñski, W., Bula, Z., and Cwojdzinski, S., 2001, U-Pb SHRIMP data confirm the presence of a Vendian foreland flysch basin next to the East European Craton: EUROPROBE Conference, Ankara, Turkey, Abstracts, p. 98–100. Zincenco, D., 1995, Chronostratigraphic scale of the pre-Permian metamorphites and granitoids from Romanian Carpathians: Geological Society of Greece, Special Publications, v. 4, no. 2, p. 647–652. Zonenshain, L.P., Kuzmin, M.I., and Natapov, L.M., 1990, Geology of the USSR: A plate-tectonic synthesis, ed. B.M. Page: Washington, D.C., American Geophysical Union, Geodynamic Series 21, 242 p. Zulauf, G., Bues, C., Dörr, W., and Vejnar, Z., 2002, 10km minimum throw along the West Bohemian shear zone: Evidence for dramatic crustal thickening and high topography in the Bohemian Massif (European Variscides): International Journal of Earth Sciences, v. 91, p. 850–864, doi: 10.1007/s00531-001-0250-y. Z˙ ylin´ska, A., 2002, Stratigraphic and biogeographic significance of Late Cambrian trilobites from Lysogóry (Holy Cross Mountains, central Poland): Acta Geologica Polonica, v. 52, p. 217–238. MANUSCRIPT ACCEPTED BY THE SOCIETY 3 OCTOBER 2006
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Geological Society of America Special Paper 423 2007
Crete and the Minoan terranes: Age constraints from U-Pb dating of detrital zircons G. Zulauf* S.S. Romano W. Dörr Institut für Geowissenschaften, Universität Frankfurt a.M., Altenhöferallee 1, 60054 Frankfurt a.M., Germany J. Fiala Geologický ústav, Akademie Veˇd Cˇ eské Republiky, 16500 Praha 6, Suchdol, Czech Republic
ABSTRACT U-Pb dating of detrital zircons separated from paragneiss of the Myrsini Crystalline Complex of eastern Crete yielded peaks in 207Pb/206Pb ages at 0.6, 0.8, 1.0, 2.0, and 2.5 Ga. A striking Mesoproterozoic age gap is present between 1.1 and 1.6 Ga. The new data are compatible with U-Pb zircon ages derived from surrounding crystalline complexes of the Cyclades, the Menderes Massif, Egypt, and the Levant. Possible provenances of the zircons of the eastern Mediterranean domains are the Sahara metacraton, the Arabian-Nubian Shield, and the Kibaran belt of central Africa. Because the age spectra of the eastern Mediterranean crystalline complexes differ significantly from those of the Cadomian- and Avalonian-type terranes, they are regarded as a separate collection of peri-Gondwanan terranes referred to as Minoan terranes. In late Neoproterozoic to ?Cambrian times, the latter underwent Andean-type orogeny at the northern border of East Gondwana, close to Egypt and the Levant. There is no evidence that the Minoan terranes traveled for long distances in Phanerozoic times. Keywords: Crete, Minoan terranes, peri-Gondwanan terranes, U-Pb-TIMS, Gondwana, detrital zircons INTRODUCTION There is considerable disagreement on the location and configuration of peri-Gondwanan terranes in Neoproterozoic and Cambrian times (e.g., Murphy et al., 2004). This holds particularly for basement complexes of the eastern Mediterranean, the Neoproterozoic to Cambrian evolution of which is only poorly constrained because of strong Variscan (Hercynian) and Alpine *E-mail:
[email protected].
deformation and metamorphism. Pre-Alpine basement in the eastern Mediterranean is exposed in the Strandja Massif of Bulgaria (Okay et al., 2001), the Sakarya zone (Okay, 2000), and the Pontides (Ustaömer, 1999; Ustaömer and Rogers, 1999; Chen et al., 2002). These basement units are located north of the VardarIzmir-Ankara Suture. South of this suture, pre-Alpine basement has been found within the Menderes Massif (Bozkurt and Oberhänsli, 2001), Kithira (Xypolias et al., 2005), the Cyclades (e.g., Henjes-Kunst and Kreuzer, 1982), the Dodecanes (Franz et al., 2005), and Crete (Seidel et al., 1982). Some of these basement
Zulauf, G., Romano, S.S., Dörr, W., and Fiala, J., 2007, Crete and the Minoan terranes: Age constraints from U-Pb dating of detrital zircons, 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. 401–411, doi: 10.1130/2007.2423(19). For permission to copy, contact
[email protected]. ©2007 Geological Society of America. All rights reserved.
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complexes include ca. 1 Ga zircons that do not occur in Cadomian-type terranes of central Europe (e.g., Armorica, SaxoThuringia, Tepla-Barrandian) but are widespread in Avaloniantype terranes (see Fig. 9 in Linnemann et al., 2004, and references therein). The 1 Ga zircons of Avalonia have been attributed to the Grenvillian orogeny and are used to infer an Amazonian provenance (Nance and Murphy, 1994; Murphy et al., 2004). Thus, the question arises whether the basement complexes of the eastern Mediterranean also underwent Late Neoproterozoic orogeny at the northern margin of West Gondwana, close to the Amazonian craton, and were subsequently shifted to their present position by dextral strike-slip. The latter option has been suggested by Winchester et al. (2002). Another source for 1 Ga zircons, however, could be the Arabian-Nubian Shield or other places in northern or central Africa, as is indicated by the zircon ages of the Cambrian Elat sandstones of southern Israel (Avigad et al., 2003; Fig. 1A). The present study is focusing on the pre-Alpine basement of Crete, which is characterized by Cambrian plutonism (Romano et
CambrianOrdovician sandstone
Cratonic basement
Fault Subduction zone
Black Sea
Aegaean Sea Menderes Massif
N
Pontides
Sakarya zone
Remobilized during pan-African orogeny
Paleozoic: younger than Ordivician
Pan-African juvenile crust
ja nd f ra si St as M
Variscan orogens and remobilized older basement Mesozoic including variscan basement
MesozoicCenozoic
al., 2004) and Carboniferous to Triassic orogenic imprints (Seidel et al., 1982; Finger et al., 2002; Romano et al., 2006). The age of the country rocks of the Cambrian plutons has not yet been constrained, and the paleogeographic position of Crete during the Cadomian (Pan-African) orogenic cycle is still an open question. Does Crete belong to the so-called peri-Gondwanan terranes (Stampfli, 2000; Murphy et al., 2004), which underwent Neoproterozoic Andeantype deformation and metamorphism at the northern border of Gondwana? To answer this question and to constrain the paleogeographic position of Crete during Neoproterozoic or Cambrian times, age spectra of detrital zircons of metasediments will be used. In the present paper we present new 207Pb-206Pb single-zircon ages of the Sfaká Paragneiss, which is part of the pre-Alpine basement of eastern Crete. The spectra of these detrital zircons reflect magmatic and metamorphic imprints in different source regions. The new data will be compared with zircon ages obtained from Gondwana, from pre-Alpine basement complexes of the eastern Mediterranean ream, and from further peri-Gondwanan terranes.
B
Cycladic Blueschist Unit
B
A
Kirsehir block
Cretan Sea Cretan basement
100 km
A
300 km
Mediterranean Sea
Atlantic Ocean
Elat sandstones
Hoggar Massif
Arabian-Nubian Shield Sahara Metacraton
W
es
tA
fri c
an
Cr
at
on
Afif terrane
Figure 1. Geological map of (A) northern Africa and (B) the eastern Mediterranean (after Stern, 2002, and Romano et al., 2004).
Crete and the Minoan terranes Geological Setting Crete consists of several tectonic nappes that have been stacked together during Cenozoic subduction and collision. The lower nappes consist of the Permian to Oligocene Plattenkalk Unit, the Triassic to Lower Jurassic Tripali Unit, and the Carboniferous to Triassic Phyllite-Quartzite Unit, all of which have been affected by Alpine high-pressure (P), low-temperature (T) metamorphism and related deformation (e.g., Seidel et al., 1982; Theye et al., 1992; Fassoulas, 1998; Zulauf et al., 2002). The upper nappes include the Triassic to Eocene Tripolitza and Pindos Units as well as the “uppermost” ophiolite- and mélange-bearing unit. The Pindos nappe is free from Alpine metamorphism, whereas the Tripolitza nappe shows evidence of very low-grade metamorphism (Feldhoff et al., 1991; Klein et al., 2004; Rahl et al., 2005). The pre-Alpine crystalline basement of eastern Crete forms the middle part of the Phyllite-Quartzite Unit (Fig. 2). It is well exposed in the ENE-WSW-trending Myrsini syncline north of the Orno Oros (Fig. 3). The basement rocks have been thrust over upper Carboniferous to Triassic marbles and phyllites (Krahl et al., 1986; Zulauf et al., 2002). On top of the basement, a Triassic variegated sequence is present that consists of metavolcanics and -sediments. The pre-Alpine basement can be subdivided into four separate units that differ in age and grade of Barrovian-type metamorphism: the Vai crystalline basement, the Kalavros crystalline basement, the Myrsini crystalline basement, and the Chamezi crystalline basement (Franz et al., 2005; Romano et al., 2006). The Kalavros crystalline basement and Myrsini crystalline basement consist of mica schist, paragneiss, orthogneiss, amphibolite, quartzite, and marble that underwent amphibolite-facies metamorphism at T = 580–630 °C and P = 6.5–8.0 kilobars (Franz et al., 2005). U-Th-Pb electron microprobe dating of metamorphic monazite of gneiss and mica schist yielded Carboniferous ages (ca. 330 Ma) for the metamorphic overprint in the Myrsini crystalline basement and Permian ages (ca. 260 Ma) for the metamorphism in the Kalavros crystalline basement (Finger et al., 2002; Romano et al., 2006). K-Ar dating of muscovite and hornblende yielded cooling ages of between 310 and 290 Ma. Many muscovites and hornblendes were affected by Alpine argon loss that caused significant age spread between 204 and 314 Ma (Seidel et al., 1982). The Chamezi crystalline basement consists of mica schist, paragneiss, and orthogneiss, which show evidence of upper greenschist-facies metamorphism at T = 500–550 °C and P = 5.5–6.5 kilobars (Franz et al., 2005). Jurassic exhumation and uplift of the Myrsini crystalline basement, Chamezi crystalline basement, and Vai crystalline basement to higher structural levels is indicated by the Jurassic fission-track ages of zircon (Romano et al., 2006). The degree of Alpine metamorphism of the Phyllite-Quartzite Unit of eastern Crete has been roughly determined at 300 ± 50 °C and 800 ± 300 MPa (Theye et al., 1992). In the mica schist and gneiss of the basement, saussuritization of feldspar,
403
Protolith age
Lithology (schematic)
U
M M M M P
M
M M
L
Figure 2. Simplified tectonostratigraphic scheme of the PhylliteQuartzite Unit of eastern Crete. Sources of stratigraphic ages: Krahl et al. (1986); Kozur and Krahl (1987); Haude (1989).
albitization of plagioclase, and chloritization of biotite, staurolite, and garnet are attributed to high levels of fluid activity during Alpine deformation (Zulauf et al., 2002). Due to this fluid impact, the zircons of the basement complexes display disturbed zonation and solution, resulting in strong radiogenic Pb loss, Ca and Al gain, and Zr depletion (Romano et al., 2004). Analytical Procedure After the removal of weathered surfaces, the samples were crushed. Zircons were separated using a Wilfley table, density fractionation, and a Frantz isodynamic separator. The separation of the density fraction was carried out at the Geological Institute at the Czech Academy of Science, Prague. Minerals were handpicked under a binocular.
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U U U A M U “flysch”
Figure 3. Geological map of eastern Crete (after Zulauf et al., 2002). Black star indicates sample location of the dated Sfaká Paragneiss.
Crete and the Minoan terranes Zircons were solved and spiked with a 235U-205Pb tracer and afterward Pb and U separation was accomplished using ion-exchange resin (Bio-Rad, Ag 1 × 8, 100–200 mesh), following the system of Krogh (1982). The separated U and Pb were loaded on single rhenium filaments, and the isotopic composition was measured by Thermion mass spectrometer analysis (TIMS) (see Dörr et al., 2002). The beam intensities of 207Pb and 206Pb allowed static-mode measurements with 2σ errors usually less than 1‰. The measured Pb isotopic ratios were corrected for mass fractionation (1 ± 0.3‰ per atomic mass unit), spike, blank, and initial Pb. Pb blanks were ~3–10 pg. The data processing was carried out using the software PbDat (Ludwig, 1991) and IsoPlot 2.A (Ludwig, 2007). Ages were calculated using the decay constants of Jaffey et al. (1971). The initial Pb ratios were taken from Stacey and Kramers (1975). RESULTS To date the zircons of the Sfaká Paragneiss, ~50 kg of sample material was collected from an exposure along a path cut ~1 km north of Sfaká village (Fig. 3). The felsic gneiss is intercalated within mica schists of the Myrsini crystalline basement. The following minerals have been found: quartz, plagioclase, biotite, white mica (sericite), and opaque phases. The penetrative foliation of the gneiss results from retrograde mylonitic shearing. There is a shape-preferred orientation of stretched quartz and plagioclase, both of which are recrystallized and form porphyroclasts within a fine-grained mylonitic matrix. The plagioclases also show fracturing and strong alteration to sericite and other secondary phases. Similar to the Cretan orthogneiss samples described by Romano et al. (2004, 2006), the zircons are discordant and many of them show incorporation of common Pb, the latter resulting in low 206Pb/204Pb ratios. Because the investigated sample contained multitudes of zircons, only single zircon grains larger than 125 µm were analyzed. The shape, color, and weight of the zircons are listed in Table 1. Most of the zircons were milky pink, colorless, yellow, or purple and were rounded to subangular or euhedral. There are no striking differences in Pb loss and U content between euhedral, subangular, and rounded zircons. The maximum and minimum apparent 207Pb/206Pb ages obtained are 2823 ± 2 and 249 ± 41 Ma, respectively. The distribution of 207Pb/206Pb ages shows major peaks at ca. 0.6, ca. 0.8, and ca. 1.0 Ga. Moderate peaks occur at ca. 2.0 and ca. 2.5 Ga (Table 1, Fig. 4). There is a striking Mesoproterozoic age gap between 1.1 and 1.6 Ga. Most of the ca. 1.0 Ga zircons are rounded to subangular. Euhedral zircons largely yielded ages of ca. 2.0 Ga or even greater. DISCUSSION The rounded to subangular shape of most of the zircons and the large spread in zircon ages are consistent with our assumption that the investigated Sfaká Gneiss of the Myrsini crystalline
405
basement is actually a paragneiss. All of the thirty-seven zircons separated are discordant. The youngest zircon has a 207Pb-206Pb age of 249 ± 41 Ma. An additional Phanerozoic age has been determined at 358 ± 13 Ma. The remaining ages are Precambrian. Because the impact of Alpine subduction-related fluids on the zircons was strong (Romano et al., 2004), we interpret the Phanerozoic ages to have resulted from loss of radiogenic Pb due to Alpine fluid migration. The Precambrian ages of the detrital zircons, however, might reflect magmatic or metamorphic events in different provenances. The depositional age of the paragneiss protolith is only poorly constrained. The age of the metamorphism of the Myrsini crystalline basement rocks has been determined at ca. 330 Ma (U-Th-Pb electron microprobe dating of monazite, Finger et al., 2002), meaning that the protolith of the Sfaká Paragneiss should be older than this age. A further age constraint is the Cambrian protolith of an orthogneiss sampled from the Myrsini crystalline basement (514 ± 14 Ma, U-Pb on zircon, Romano et al., 2004). The protolith of the Sfaká Paragneiss should belong to the country rocks of this granitoid protolith, and the deposition of the paragneiss protolith should be older than Upper Cambrian (Furongian). An upper limit for the age of the paragneiss protolith is given by the large number of Late Neoproterozoic detritic zircons (547 ± 27, 556 ± 18, 573 ± 32, 592 ± 15, 596 ± 26, and 603 ± 14 Ma). Taking into account the numbers mentioned earlier, the protolith age of the Sfaká Paragneiss is bracketed by the period from Latest Neoproterozoic to Upper Cambrian. The assumption that the Precambrian ages of the detrital zircons might reflect magmatic or metamorphic events in different source areas is supported by the fact that the Cretan zircon ages are strikingly similar to zircon ages obtained from surrounding crystalline complexes. This holds particularly for the MenderesTaurus block, whose zircons show approximately the same age pattern as those of the Sfaká Paragneiss (Fig. 5C). Apart from a few Archean 207Pb-206Pb ages, the Menderes-Taurus block is characterized by zircon ages occupying the entire Neoproterozoic period. Further age clusters occur at ca. 1.9 and ca. 2.5 Ga, whereas Mesoproterozoic ages are lacking (Kröner and Şengör, 1990; Hetzel et al., 1998; Loos and Reischmann, 1999; Gessner et al., 2004). Keay et al. (2001) and Keay and Lister (2002) investigated zircons of the Cycladic gneisses using a sensitive high-resolution ion microprobe. The distribution of Precambrian detrital and inherited zircon ages is again similar to that described earlier (Fig. 5A). There are Neoproterozoic ages with a major peak at ca. 0.6 Ga and a moderate peak at 0.9 Ga. Moderate Paleoproterozoic age maxima are present at ca. 2.0 and ca. 2.4 Ga. Mesoproterozoic ages are conspicuously scarce. The detrital zircons of a Cambrian sandstone of Elat, southern Israel, show patterns partly similar to those described earlier (Avigad et al., 2003; Fig. 5B). This holds particularly for the Neoproterozoic period. Neoproterozoic ages are widespread, with a distinct peak at ca. 0.6 Ga. Similar to the age patterns of the Cyclades, the zircon ages of the Elat sandstone also covering the
ro, r r, kl, e m, kl, r, lp ro, r r, ro, m ro, r ro, r r, lp, ro k, ro ro, k ro, m, k, lp ro, k, lp r, k f, m, k ro, m, k r, ro, k, y k, kl, ro ro, kl, k ro, lp, k ro, k, k ro, k, r ro, kl ro, kl ro f, kl, e ro, k lp, r, ro, m ro, kl lp, ro, kl, k p, g, k p, m ro, k, kl, lp ro, m, k-i p, m
r, ro, e f, m, r
Typology
Weight μ (g) 20 17 28 10 17 27 16 10 13 17 19 37 10 10 13 12 17 20 14 15 11 19 17 13 12 10 13 12 12 27 15 40 12 10 18 15 15
43.1 11.2 22.3 74.6 10.4 14.3 12.9 117.7 70.6 54.2 151.8 13.7 39.4 24.6 19.0 37.9 11.4 7.8 9.5 115.9 37.6 29.4 31.9 33.1 9.9 13.2 36.2 26.8 30.5 21.2 84.4 30.4 50.0 158.2 50.8 166.0 38.5
Pb r. (ppm) 3.8 1.6 3.4 10.7 0.8 2.8 4.1 2.0 4.2 1.4 1.1 3.1 12.2 8.2 4.0 5.7 6.1 1.1 0.5 0.4 4.7 2.8 11.1 18.0 2.9 2.0 2.2 0.9 3.6 1.2 0.6 0.2 2.0 7.7 4.0 2.4 15.6
Pb i. (ppm)
198 443 304 204 464 100 123 124 739 236 198 196 190 331 215 441 205 191 154 254 92 327 1314 136 392 115
864 141 224 1900 86 121 129 689 265 161
U (ppm)
TABLE 1. U-Pb DATA FROM ZIRCONS OF THE SFAKÁ PARAGNEISS Radiogenic ratios Pb/204Pb* 207 207 206 Pb/235U Pb/238U Pb/206 Pb 2σ 2σ 723 0.04985 0.28 0.3690 0.65 0.05370 452 0.07979 0.83 0.6509 1.72 0.05917 382 0.08781 0.31 0.7226 0.76 0.05968 444 0.03875 0.37 0.3248 0.72 0.05999 678 0.10054 1.13 0.9459 1.94 0.06824 288 0.10138 0.47 0.9783 0.95 0.06999 210 0.10067 0.80 0.9758 1.30 0.07031 3079 0.14853 0.33 1.4737 0.52 0.07196 837 0.23310 0.29 4.8380 0.36 0.15053 1757 0.28073 0.33 6.7223 0.35 0.17368 6978 0.19962 278 0.06528 0.35 0.5282 0.91 0.05870 203 0.08275 0.50 0.6823 1.32 0.05980 185 0.07332 0.69 0.6513 1.60 0.06443 303 0.09103 0.73 0.8097 1.34 0.06451 437 0.08420 0.46 0.7579 0.75 0.06529 112 0.09419 0.90 0.8501 2.10 0.06546 420 0.05867 1.11 0.5394 2.79 0.06668 1010 0.07177 1.29 0.6755 1.60 0.06827 18,106 0.15191 0.17 1.5205 0.20 0.07260 443 0.13788 0.54 1.3933 0.81 0.07329 599 0.13792 0.35 1.4102 0.52 0.07416 174 0.14751 0.34 1.5114 1.04 0.07432 115 0.16367 0.70 2.5679 1.39 0.11379 244 0.03116 1.24 0.2199 2.23 0.05119 418 0.05901 1.25 0.4755 1.79 0.05845 1068 0.08477 0.86 0.7694 3.40 0.06583 1719 0.12867 0.52 1.2953 0.75 0.07301 438 0.13680 0.52 1.9333 0.68 0.10250 969 0.12452 0.43 1.9992 0.62 0.11645 6337 0.26276 0.21 4.3425 0.234 0.11986 8897 0.29742 0.27 .4933 0.31 0.12030 1140 0.11126 0.60 1.8839 0.85 0.12280 1191 0.12172 0.17 2.1930 0.21 0.13067 646 0.33035 0.24 6.7416 0.29 0.14801 3594 0.39998 0.53 9.2437 0.55 0.16762 111 0.23935 0.44 5.5796 0.91 0.16907 206
2σ 0.58 1.48 0.68 0.65 1.56 0.80 0.98 0.40 0.22 0.12 0.12 0.81 1.19 1.47 1.10 0.57 1.81 2.53 0.91 0.11 0.59 0.37 0.94 1.14 1.78 1.24 3.29 0.53 0.43 0.45 0.10 0.16 0.59 0.11 0.15 0.15 0.75 0.46 0.45 0.47 0.57 0.65 0.51 0.42 0.82 0.84 0.69 0.70 0.45 0.58 0.60 0.72 0.26 0.71 0.78 0.69 0.91 0.87 0.72 0.86 0.85 0.96 0.58
0.46 0.50 0.46 0.42 0.60 0.54 0.66 0.65 0.80 0.94
Rho
408 512 456 561 521 580 368 447 912 832 833 886 977 197 369 525 780 826 757 1504 1679 680 740 1840 2168 1383
Pb/ 238U 314 494 542 245 618 622 618 893 1350 1595
206
1 3 3 4 2 5 4 6 2 5 3 3 7 2 5 5 4 4 3 3 5 4 1 4 11 6
2σ 1 4 2 1 7 3 5 3 4 5 431 528 509 602 573 625 438 524 939 886 893 935 1292 202 395 579 844 1093 1115 1702 1808 1075 1179 2078 2363 1913
4 7 8 8 4 13 12 8 2 7 5 10 18 5 7 20 6 7 7 4 6 9 2 6 13 17
Age (Ma) Pb/ 235U 2σ 319 2 509 9 552 4 282 2 676 13 693 7 691 9 920 5 1792 6 2076 7 207
Pb/206 Pb 358 573 592 603 876 928 937 985 2352 2593 2823 556 596 756 759 784 789 828 877 1003 1022 1046 1050 1861 249 547 801 1014 1670 1902 1954 1961 1997 2107 2323 2534 2548
207
2σ 13 32 15 14 32 17 20 8 4 2 2 18 26 31 23 12 38 53 19 2 12 8 19 21 41 27 69 11 8 8 2 3 11 2 3 3 13
Note: Errors are quoted at the 2 sigma (2σ) level. Abbreviations: Pb r.—radiogenic lead; Pb i.—initial common lead; e—inclusions; f—colorless; g—stumpy; k—subangular; kl—clear; lp—long prismatic; m—milky; p—purple; r—rounded; ro—pink; y—yellow. *Corrected for mass fraction (1.12 ± 0.18‰ per atomic mass unit) and analytical blank.
1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22 23 24 25 26 27 28 29 30 31 32 33 34 35 36 37
Sample
406 Zulauf et al.
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Number of analyses
Crete and the Minoan terranes
Age (GA) Figure 4. Histogram showing different Precambrian peaks in 207Pb/206Pb ages of Sfaká Paragneiss zircons.
Neoproterozoic-Mesoproterozoic boundary. However, there is also a lack in Mesoproterozoic ages between ca. 1.3 and 1.5 Ga. On the other hand, age peaks at ca. 1.9–2.0 and ca. 2.4–2.5 Ga do not occur (Avigad et al., 2003). Because all of the eastern Mediterranean areas considered are located at or close to the northern border of Afro-Arabia, it is most likely that this area was the area of provenance of the zircons. Three main orogenic cycles in Africa overprinted the margins of older cratons (Kröner, 1977; Cahen et al., 1984): Eburnian (2200–1800 Ma), Kibaran (1350–950 Ma), and PanAfrican (800–500 Ma). West Africa is characterized by ages of 1.8–2.2 Ga and 2.8–3.4 Ga (e.g., Nance and Murphy, 1994). Pan-African and Eburnian ages are widespread in the Saharan metacraton (Abdelsalam et al., 2002; Fig. 5D). Ages at ca. 2.1 Ga have also been determined from rocks of the Afif terrane of the Arabian-Nubian Shield (Stacey and Hedge, 1984). The Saharan metacraton further shows a distinct age peak at the ProterozoicArchean boundary at ca. 2.4 Ga. This peak has also been found on Crete (present study), the Cyclades (Keay and Lister, 2002), and the Menderes-Taurus block (Kröner and Şengör, 1990). We therefore conclude that in Cambrian or Neoproterozoic times the Saharan metacraton was the area of provenance of Pan-African and Eburnian zircons and of zircons with ages of ca. 2.4 Ga, all of which were deposited in the eastern Mediterranean crystalline complexes mentioned earlier. Moreover, the few Mesoproterozoic ages found in the Saharan metacraton (Abdelsalam et al., 2002) could explain the occurrence of few Mesoproterozoic ages in the eastern Mediterranean domains. The oldest zircon of the Sfaká Paragneiss (2832 ± 2 Ma) is probably derived from the Hoggar Massif of northern Africa, where Archean gneisses are widespread (Peucat et al., 2003). Grenvillian and Kibaran ages, respectively, particularly those around 1.0 Ga, are lacking in the West African craton (e.g., Nance and Murphy, 1994; Ennih and Liégeois, 2001; Egal
et al., 2002) and in the Sahara metacraton (Abdelsalam et al., 2002). Grenvillian or Kibaran ages, however, are common on Crete and in the surrounding Mediterranean domains mentioned earlier. Thus there should be a further source area from which the Grenvillian and Kibaran zircons were derived. The Kibaran cycle is named after the central African Kibaran belt, the latter straddling the Democratic Republic of Congo, Uganda, Burundi, Rwanda, and Tanzania. The Kibaran orogenic belt is the result of the convergence of Paleoproterozoic and Archaean cratonic blocks (the Congo, Kalahari, Bangweulu, Tanzania, and West Nilian cratons). New geochronologic data and interpretations suggest convergence between the Kalahari craton and a composite Congo-Laurentia craton during the assembly of Rodinia, generating the Kibaran, Grenville, and Llano belts (Kampunzu et al., 2003; Fig. 6). The ages of magmatic rocks in the Kibaran orogen range from ca. 1.4 to ca. 1.0 Ga. Because the Kibaran zircon ages are present not only in the basement of Crete, the Menderes, and the Cyclades, but also in the Elat sandstones of Israel, there is no reason to say that the basement complexes of the eastern Mediterranean have traveled for long distances since Neoproterozoic times (cf. Winchester et al., 2002). However, the Kibaran basement rocks are currently exposed more than 3000 km south of Crete. If the ca. 1 Ga zircons of the eastern Mediterranean basement complexes are derived from the Kibaran orogen, the zircons must have traveled for a long distance. This assumption is compatible with our observation that most of the Kibaran zircons of the Sfaká Paragneiss are rounded or subangular, whereas most of the euhedral zircons yielded ages of around 2 Ga, pointing to a closer source in northern Africa (the Sahara metacraton). On the other hand, the possibility that the Arabian-Nubian Shield contains rocks with ca. 1 Ga zircons should also be explored (Avigad et al., 2003). Further evidence of provenance of the zircons of the Sfaká Gneiss from the Arabian-Nubian Shield is given by the large number of zircon ages
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Phaneroz. Neoprot. Mesoprot.
Palaeoproteroz.
Archaic
A
Inherited ages
20
Cyclades
Detritic ages
Keay et al., 2001 Keay and Lister, 2002
10
B
206Pb/238U
& 207Pb/206Pb concordant ages
20 207Pb/206Pb discordant ages
Elat sandstone, Israel Avigad et al., 2003
Number of analyses
5
C
8
Menderes Massif Kröner and Sengör, 1990
4
D 8
Sahara metacraton Ref. in Abdelsalam et al., 2002
4
E Cadomia
8
Fernández-Suárez et al., 2002
4
0
0.5
1
1.5 2 Age (Ga)
2.5
3
3.5
Figure 5. Published zircon ages from the eastern Mediterranean and northern Africa: (A) Cycladic metamorphic rocks (Keay et al., 2001; Keay and Lister, 2002); (B) Cambrian or Ordovican sandstone of Elat, Israel (Avigad et al., 2003); (C) Menderes-Taurus block (Kröner and Şengör, 1990); (D) Sahara metacraton (Abdelsalam et al., 2002, and references therein); and (E) Cadomia (Fernández-Suárez et al., 2002).
between 750 and 900 Ma. During this period a large volume of island arc rocks developed in the Arabian-Nubian Shield (e.g., Stern, 1994; Blasband et al., 2000, and references therein). There is a striking difference in age patterns between the Avalonian and Cadomian terranes and the crystalline complexes of the eastern Mediterranean. The Avalonian terranes (West Avalonia, East Avalonia, Carolina, Moravo-Silesia, Oaxaquia, and the Chortis block) resulted from ca. 1.3–1.0 Ga juvenile crust within the Panthalassa-type ocean surrounding Rodinia and were accreted to the northern margin of Gondwana at ca. 650 Ma (Murphy et al., 2004, and references therein). The Cadomian terranes (Armorica, Saxo-Thuringia, TeplaBarrandian, and Ossa Morena) originated along the West African margin from recycled ancient (2–3 Ga) West African crust. Apart from Neoproterozoic or Cambrian tectonometamorphic imprints, these terranes are characterized by Eburnian zircons, the latter derived from basement rocks of northwestern Africa (e.g., Piton, 1985; Dörr et al., 2002; Fernández-Suárez et al., 2002; Gutiérrez Alonso et al., 2003; Linnemann et al., 2004). Grenvillian or Kibaran and Mesoproterozoic ages, on the other hand, are lacking in all of the Cadomian terranes except in the Ossa-Morena zone (Fernández-Suárez et al., 2002; Gutiérrez Alonso et al., 2003). Because of the striking differences between the Cadomian and Avalonian terranes and the basement complexes of the eastern Mediterranean, the latter are referred to as Minoan terranes. They take their name from Minos, who, according to Greek mythology, was the king of Crete and the son of Zeus and Europa. The Minoan terranes represent the eastern extremity of peri-Gondwanan terranes (Fig. 6) that underwent Late Neoproterozoic to Cambrian Andean-type orogeny at the northern border of Gondwana (see also Robertson et al., 1996; Kusky et al., 2003; Linnemann et al., 2004). There is no evidence that the Minoan terranes traveled for long distances in Phanerozoic times. An open question concerns the end of the Cadomian (PanAfrican) cycle in this area. Late Neoproterozoic to early Cambrian magmatic activity is documented in Jordan (ca. 540 Ma, Jarrar et al., 2003), Egypt (ca. 560–780 Ma, Wilde and Youssef, 2002), the Pontides (ca. 590–560 Ma, Ustaömer, 1999; Ustaömer and Rogers, 1999; Chen et al., 2002) and the Menderes Massif (ca. 540–570 Ma, Reischmann et al., 1991; Hetzel and Reischmann, 1996; Hetzel et al., 1998; Gessner et al., 2004). Moreover, recent U-Pb dating of garnet of the Menderes Massif yielded evidence of Late Neoproterozoic metamorphism (Ring et al., 2004). The Cambrian plutons of eastern Crete are unambiguously younger (511 ± 16 and 514 ± 14 Ma, Romano et al., 2004) than the granitoids mentioned earlier. The geodynamic setting of these granitoids, however, is still an open question. The metagranitoids do not yield unequivocal geochemical signatures that could be used to constrain the geodynamic setting. In geochemical discriminating diagrams, they plot at the boundary between volcanic-arc and within-plate granitoids (Romano et al., 2006). Because the uncertainty of the 207Pb-206Pb ages given earlier is relatively large and only one type of paragneiss has been
Crete and the Minoan terranes
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Figure 6. Paleogeographic reconstruction of Gondwana (modified after Kusky et al., 2003, and Linnemann et al., 2004). PanAfrican orogens and cratons are indicated. Based on similar zircon age spectra, Crete, the Cyclades, and the Menderes Massif are referred to as Minon peri-Gondwanan terranes. The source regions for the sedimentary deposits of the Minoan terranes are the Arabian-Nubian Shield and the Sahara metacraton.
considered to constrain the age spectra of detrital zircons, the question about the validity of the new age spectra arises. Further investigations of the age of detrital zircons from different types of paragneisses are necessary to prove the geodynamic implications concluded from the present study. ACKNOWLEDGMENTS We thank P.A. Ustaömer and A. Zelazniewicz for their constructive reviews. Financial support by Deutsche Forschungsgemeinschaft (DFG, grant Zu 73–8) is gratefully acknowledged. REFERENCES CITED Abdelsalam, M.G., Liegeois, J.-P., and Stern, R.J., 2002, The Saharan Metacraton: Journal of African Earth Sciences, v. 34, p. 119–136, doi: 10.1016/ S0899-5362(02)00013-1. Avigad, D., Kolodner, K., McWilliams, M., Persing, H., and Weissbrod, T., 2003, Origin of northern Gondwana Cambrian sandstone revealed by detrital zircon SHRIMP dating: Geology, v. 31, p. 227–230, doi: 10.1130/00917613(2003)031<0227:OONGCS>2.0.CO;2.
Blasband, B., White, S., Brooijmans, P., De Boorder, H., and Visser, W., 2000, Late Proterozoic extensional collapse in the Arabian-Nubian Shield: Journal Geological Society London, v. 157, p. 615–628. Bozkurt, E., and Oberhänsli, R., 2001, Menderes Massif (Western Turkey): Structural, metamorphic and magmatic evolution—A synthesis: International Journal of Earth Sciences, v. 89, p. 679–708, doi: 10.1007/ s005310000173. Cahen, L., Snelling, N.J., Delhal, J., Vail, J.R., Bonhomme, M., and Ledent, D., 1984, The geochronology and evolution of Africa: Oxford, England, Clarendon, 512 p. Chen, F., Siebel, W., Satir, M., Terzioglu, M., and Saka, K., 2002, Geochronology of the Karadere basement (NW Turkey) and implications for the geological evolution of the Istanbul zone: International Journal of Earth Sciences, v. 91, p. 469–481, doi: 10.1007/s00531-001-0239-6. Dörr, W., Zulauf, G., Fiala, J., Franke, W., and Vejnar, Z., 2002, Neoproterozoic to Early Cambrian history of an active plate margin in the Teplá-Barrandian unit—A correlation of U-Pb isotopic-dilution-TIMS ages (Bohemia, Czech Republic): Tectonophysics, v. 352, p. 65–85, doi: 10.1016/S00401951(02)00189-0. Egal, E., Thiéblemont, D., Lahondére, D., Guerrot, C., Costea, C.A., Iliescu, D., Delor, C., Goujou, J.C., Lafon, J.M., Tegyey, M., Diaby, S., and Kolié, P., 2002, Late Eburnean granitization and tectonics along the western and northwestern margin of the Archean Kénéma-Man domain (Guinea, West African Craton): Precambrian Research, v. 117, p. 57–84, doi: 10.1016/ S0301-9268(02)00060-8.
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Ennih, N., and Liégeois, J.P., 2001, The Moroccan Anti-Atlas: The West African craton passive margin limited Pan-African activity. Implications for the northern limit of the craton: Precambrian Research, v. 112, p. 289– 302, doi: 10.1016/S0301-9268(01)00195-4. Fassoulas, C., 1998, The structural evolution of central Crete: Insight into the tectonic evolution of the south Aegean (Greece): Journal of Geodynamics, v. 27, p. 23–45, doi: 10.1016/S0264-3707(97)00026-4. Feldhoff, R.A., Luecke, A., and Richter, D.K., 1991, Ueber die Diagenese/ Metamorphosebedingungen der Pindos- und Tripolitza-Serie auf der Insel Kreta (Griechenland): Zentralblatt für Geologie und Palaeontologie, v. 1990, no. 11, p. 1611–1622. Fernández-Suárez, J., Gutiérrez-Alonso, G., and Jeffries, T.E., 2002, The importance of along-margin terrane transport in northern Gondwana: Insights from detrital zircon parentage in Neoproterozoic rocks from Iberia and Brittany: Earth and Planetary Science Letters, v. 204, p. 75–88, doi: 10.1016/S0012-821X(02)00963-9. Finger, F., Krenn, E., Riegler, G., Romano, S., and Zulauf, G., 2002, Resolving Cambrian, Carboniferous, Permian and Alpine monazite generations in the polymetamorphic basement of eastern Crete (Greece) by means of the electron microprobe: Terra Nova, v. 14, p. 233–240, doi: 10.1046/j.13653121.2002.00415.x. Franz, L., Okrusch, M., Seidel, E., and Kreuzer, H., 2005, Polymetamorphic evolution of pre-Alpidic basement relics in the external Hellenides, Greece: Neues Jahrbuch für Mineralogie-Abhandlungen, v. 181, p. 147– 172, doi: 10.1127/0077-7757/2005/0013. Gessner, K., Collins, A.S., Ring, U., and Güngör, T., 2004, Resolving the structural and thermal history of poly-orogenic basement: SHRIMP geochronology of granitoid rocks in the southern Anatolide belt, western Turkey: Journal of the Geological Society, London, v. 161, p. 93–101. Gutiérrez-Alonso, G., Fernández-Suárez, J., Jeffries, T.E., Jenner, G.A., Tubrett, M.N., Cox, R., and Jackson, S.E., 2003, Terrane accretion and dispersal in the northern Gondwana margin: An Early Paleozoic analogue of a long-lived active margin: Tectonophysics, v. 365, p. 221–232, doi: 10.1016/S0040-1951(03)00023-4. Haude, G., 1989, Geologie der Phyllit-Einheit im Gebiet um Palekastro (NordostKreta, Griechenland) [Ph.D. thesis]: Munich, Technische Universität. Henjes-Kunst, F., and Kreuzer, H., 1982, Isotopic dating of pre-Alpidic rocks from the island of Ios (Cyclades, Greece): Contributions to Mineralogy and Petrology, v. 80, p. 245–253, doi: 10.1007/BF00371354. Hetzel, R., and Reischmann, T., 1996, Intrusion age of Pan-African augen gneisses in the southern Menderes massif and the age of cooling after Alpine ductile extensional deformation: Geological Magazine, v. 133, p. 565–572. Hetzel, R., Romer, R.L., Candan, O., and Passchier, C.W., 1998, Geology of the Bozdag area, central Menderes massif, SW Turkey: Pan-African basement and Alpine deformation: Geologische Rundschau, v. 87, p. 394–406, doi: 10.1007/s005310050218. Jaffey, A.H., Flynn, K.F., Glendenin, L.E., Bentley, W.C., and Essling, A.M., 1971, Precision measurements of half-lives and specific activities of 235U and 238U: Physical Review, Section C. Nuclear Physics, v. 4, p. 1889–1906. Jarrar, G., Stern, R.J., Saffarini, G., and Al-Zubi, H., 2003, Late- and post-orogenic Neoproterozoic intrusions of Jordan: Implications for crustal growth in the northernmost segment of the East African Orogene: Precambrian Research, v. 123, p. 295–319, doi: 10.1016/S0301-9268(03)00073-1. Kampunzu, A.B., Milesi, J.P., and Deschamps, Y., 2003, Africa within Rodinia supercontinent: Evidence from the Kibaran orogenic system: Boulder, Colorado, Geological Society of America Abstracts with Programs, v. 35, p. 301. Keay, S., and Lister, G.S., 2002, African provenance for the metasediments and metaigneous rocks of the Cyclades, Aegean Sea, Greece: Geology, v. 30, p. 235–238. Keay, S., Lister, G.S., and Buick, I., 2001, The timing of partial melting, Barrovian metamorphism and granite intrusion in the Naxos metamorphic core complex, Cyclades, Aegean Sea, Greece: Tectonophysics, v. 342, p. 275–312, doi: 10.1016/S0040-1951(01)00168-8. Klein, T.H., Zulauf, G., Craddock, J.P., and Heidelbach, F., 2004, Metamorphic Tripolitza rocks on top of non-metamorphic sediments: Consequences for the geological evolution of Crete: 10th International Congress of the Geological Society of Greece, Thessaloniki, Greece (abstract). Krahl, J., Kauffmann, G., Richter, D., Kozur, H., Möller, I., Förster, O., Heinritzi, F., and Dornsiepen, U., 1986, Neue Fossilfunde in der PhyllitGruppe Ostkretas (Griechenland): Zeitschrift der Deutschen Geologischen Gesellschaft, v. 137, p. 523–536.
Krogh, T.E., 1982, Improved accuracy of U-Pb zircon ages by the creation of more concordant systems using an air abrasion technique: Geochimica et Cosmochimica Acta, v. 46, p. 637–649, doi: 10.1016/00167037(82)90165-X. Kröner, A., 1977, Precambrian plate tectonics. Amsterdam, Elsevier, 801 p. Kröner, A., and Şengör, A.M.C., 1990, Archean and Proterozoic ancestry in late Precambrian to early Paleozoic crustal elements of southern Turkey as revealed by single-zircon dating: Geology, v. 18, p. 1186–1190, doi: 10.1130/0091-7613(1990)018<1186:AAPAIL>2.3.CO;2. Kusky, T.M., Abdelsalam, M.G., Stern, R.J., and Tucker, R.D., 2003, Evolution of the East African and related orogens, and the assembly of Gondwana: Precambrian Research, v. 123, p. 81–85, doi: 10.1016/S03019268(03)00062-7. Linnemann, U., McNaughton, N.J., Romer, R.L., Gehmlich, M., Drost, K., and Tonk, C., 2004, West African provenance for Saxo-Thuringia (Bohemian Massif): Did Armorica ever leave pre-Pangean Gondwana? U-Pb-SHRIMP zircon evidence and the Nd-isotopic record: International Journal of Earth Sciences, v. 93, p. 683–705, doi: 10.1007/s00531-004-0413-8. Loos, S., and Reischmann, Th., 1999, The evolution of the southern Menderes Massif in SW Turkey as revealed by zircon dating: Journal of the Geological Society, London, v. 156, p. 1021–1030. Ludwig, K.R., 1991, PBDAT version 1.23, Open File: Reston, Virginia, U.S. Geological Survey. Ludwig, K.R., 2007, Isoplot/EX: A geochronological toolkit for Microsoft Excel: Berkeley, California, Berkeley Geochronological Center, Special Publication 1a. Murphy, J.B., Pisarevsky, S.A., Nance, R.D., and Keppie, J.D., 2004, Neoproterozoic–Early Paleozoic evolution of peri-Gondwanan terranes: Implications for Laurentia-Gondwana connections: International Journal of Earth Sciences, v. 93, p. 659–682, doi: 10.1007/s00531-004-0412-9. Nance, R.D., and Murphy, J.B., 1994, Contrasting basement isotopic signatures and the palinspastic restoration of peripheral orogens: Example from the Neoproterozoic Avalonia-Cadomian belt: Geology, v. 22, p. 617–620, doi: 10.1130/0091-7613(1994)022<0617:CBISAT>2.3.CO;2. Okay, A.I., 2000, Was the Late Triassic orogeny in Turkey caused by the collision of an oceanic plateau? in Bozkurt, E., Winchester, J.A., and Piper, D.J.W., eds., Tectonics and magmatism in Turkey and the surrounding area: Geological Society, London, Special Publication 173, p. 25–42. Okay, A.I., Satir, M., Tüysüz, O., Akyüz, S., and Chen, F., 2001, The tectonics of the Strandja Massif: Late-Variscan and mid-Mesozoic deformation and metamorphism in the northern Aegean: International Journal of Earth Sciences, v. 90, p. 217–233, doi: 10.1007/s005310000104. Peucat, J.J., Drareni, A., Latouche, L., Deloule, E., and Vidal, P., 2003, U-Pb zircon (TIMS and SIMS) and Sr-Nd whole-rock geochronology of Gour Oumelalen granulitic basement, Hoggar massif, Tuareg shield, Algeria: Journal of African Earth Sciences, v. 37, p. 229–239, doi: 10.1016/j.jafre arsci.2003.03.001. Piton, P., 1985, Etude géochimique et géochronologique de la croûte continental protérozoïque (1.8–2.2) au Nord du Massif Armoricain [Raport de Diplôme d’Etudes Approfondies]: Rennes, France, Université de Rennes. Rahl, J.M., Anderson, K.M., Brandon, M.T., and Fassoulas, C., 2005, Raman spectroscopic carbonaceous material thermometry of low-grade metamorphic rocks: Calibration and application to tectonic exhumation in Crete, Greece: Earth and Planetary Science Letters, v. 240, p. 339–354. Reischmann, T., Kröner, A., Todt, W., Dürr, S., and Şengör, A.M.C., 1991, Episodes of crustal growth in the Menderes Massif, W Turkey, inferred from zircon dating: Terra Abstracts, v. 3, p. 34. Ring, U., Buchwaldt, R., and Gessner, K., 2004, Pb-Pb dating of garnet from the Anatolide belt in western Turkey: Regional implications and speculations on the role Anatolia played during the amalgamation of Gondwana: Zeitschrift deutsche geologische Gesellschaft, v. 154, p. 537–555. Robertson, A.H.F., Dixon, J.E., Brown, S., Collins, A., Morris, A., Pickett, E., Sharp, I., and Ustaömer, T., 1996, Alternative tectonic models for the Late Palaeozoic–Early Tertiary development of Tethys in the Eastern Mediterranean region, in Morris, A., and Tarling, D.H., eds., Palaeomagnetism and tectonics of the Mediterranean region: Geological Society, London, Special Publication 105, p. 239–263. Romano, S.S., Dörr, W., and Zulauf, G., 2004, Cambrian granitoids in preAlpine basement of Crete (Greece): Evidence from U-Pb dating of zircon, in Dörr, W., Finger, F., Linnemann, U., et al., eds., The Avalonian-Cadomian Belt and related peri-Gondwana terranes: International Journal of Earth Sciences, v. 93, p. 844–859.
Crete and the Minoan terranes Romano, S.S., Brix, M.R., Dörr, W., Fiala, J., Krenn, E., and Zulauf, G., 2006, The Carboniferous to Jurassic evolution of the pre-Alpine basement of Crete: Constraints from radiometric dating of orthogneiss, fission-track dating of zircon and structural/petrological data, in Robertson, A.H.F., and Mountrakis, D., eds., Tectonic development of the Eastern Mediterranean Region: Boulder, Colorado, Geological Society, London, Special Publication 260, p. 69–90. Seidel, E., Kreuzer, H., and Harre, W., 1982, A Late Oligocene / Early Miocene high pressure belt in the External Hellenides: Geologisches Jahrbuch, v. E23, p. 165–206. Stacey, J.S., and Hedge, C.E., 1984, Geochronologic and isotopic evidence for early Proterozoic crust in the eastern Arabian shield: Geology, v. 12, p. 310–313, doi: 10.1130/0091-7613(1984)12<310:GAIEFE>2.0.CO;2. Stacey, J.S., and Kramers, J.D., 1975, Approximation of terrestrial lead isotope evolution by a two-stage model: Earth and Planetary Science Letters, v. 26, p. 207–221, doi: 10.1016/0012-821X(75)90088-6. Stampfli, G., 2000, Tethyian oceans, in Bozkurt, E., Winchester, J.A., and Piper, D.J.W., eds., Tectonics and magmatism in Turkey and the surrounding area: Geological Society, London, Special Publication 173, p. 1–24. Stern, R.J., 1994, Arc assembly and continental collision in the Neoproterozoic East African Orogen: Implications for the consolidation of Gondwanaland: Annual Review of Earth and Planetary Sciences, v. 22, p. 319–351. Theye, T., Seidel, E., and Vidalo, O., 1992, Carpholite, sudoite and chloritoid in low-temperature high-pressure metapelites from Crete and the Peloponnesus, Greece: European Journal of Mineralogy, v. 4, p. 487–507. Ustaömer, P.A., 1999, Pre–Early Ordovician Cadomian arc-type granitoids, the Bolu Massif, West Pontides, northern Turkey: Geochemical evidence:
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International Journal of Earth Sciences, v. 88, p. 2–12, doi: 10.1007/ s005310050241. Ustaömer, P.A., and Rogers, G., 1999, The Bolu Massif: Remnant of a pre–Early Ordovician active margin in the west Pontides, northern Turkey: Geological Magazine, v. 136, p. 579–592, doi: 10.1017/S0016756899003015. Wilde, S.A., and Youssef, K., 2002, A re-evaluation of the origin and setting of the Late Precambrian Hammamat Group based on SHRIMP U-Pb dating of detrital zircons from Gebel Umm Tawat, North Eastern Desert, Egypt: Journal of the Geological Society, London, v. 159, p. 595–604. Winchester, J.A., Pharaoh, T.C., and Verniers, J., 2002, Palaeozoic amalgamation of Central Europe: An introduction and synthesis of new results from recent geological and geophysical investigations, in Winchester, J.A., Pharaoh, T.C., and Verniers, J., eds., Palaeozoic amalgamation of Central Europe: Geological Society, London, Special Publication 201, p. 1–18. Xypolias, P., Dörr, W., and Zulauf, G., 2005, Upper Carboniferous plutonism within the pre-Alpine basement of the External Hellenides (Kithira, Greece): Evidence from U-Pb zircon dating: Journal of the Geological Society, London, v. 163, p. 539–547. Zulauf, G., Kowalczyk, G., Krahl, J., Petschick, R., and Schwanz, S., 2002, The tectonometamorphic evolution of high-pressure low-temperature metamorphic rocks of eastern Crete, Greece: Constraints from microfabrics, strain, illite crystallinity and paleostress: Journal of Structural Geology, v. 24, p. 1805–1828.
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Printed in the USA
Geological Society of America Special Paper 423 2007
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 J. Brendan Murphy* Department of Earth Sciences, St. Francis Xavier University, Antigonish, Nova Scotia, Canada, B2G 2W5
ABSTRACT The Early Silurian–Early Devonian Arisaig Group, in the Avalon terrane of Nova Scotia, consists of a thick (~1900 m) sequence of unmetamorphosed fossiliferous siliciclastic strata that unconformably overlies the 460 Ma bimodal Dunn Point Formation volcanic rocks and is unconformably overlain by basalts and red clastic rocks of the McArras Brook Formation. The Dunn Point volcanic rocks were deposited when Avalonia was a microcontinent, in a New Zealand–type arc setting ~1800 km north of Gondwana and 1700–2000 km south of Laurentia. Geochemical, Sm-Nd, and U-Pb (detrital zircon) isotopic data of all Arisaig Group strata show fundamental differences from the underlying Avalonian rocks, indicating that they were not derived from Avalonian basement. These data are instead compatible with derivation from Baltica, implying that Avalonia had accreted to Baltica by the earliest Silurian and that the Arisaig Group is part of a clastic sequence that has overstepped Appalachian-Caledonide terrane boundaries. The lack of penetrative deformation and the approximately concordant nature of the contact between the Dunn Point Formation and the Arisaig Group suggest that this portion of Avalonia was located on the trailing edge of the Avalonia plate during the collision. Regional syntheses suggest that the basin was initiated by local transtension during oblique sinistral collision between Avalonia and Baltica. An overall trend toward increasingly negative εNd values in the clastic rocks toward the top of the Arisaig Group is thought to reflect increasing input from Laurentia by the time of deposition of the Early Devonian strata. The basin also preserves evidence of loading in the Late Silurian, which is thought to reflect the development of a foreland basin and the ongoing shortening across the orogen associated with the onset of the Acadian orogeny. The unconformity between the Arisaig Group and the overlying McArras Brook Formation is the local expression of the deformation associated with
*E-mail:
[email protected]. Murphy, J.B., 2007, 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, 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. 413–435, doi: 10.1130/2007.2423(20). For permission to copy, contact
[email protected]. ©2007 Geological Society of America. All rights reserved.
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Murphy Acadian orogeny in the Antigonish Highlands. The orientation and style of Acadian deformation preserved in the Arisaig Group is compatible with dextral movement along major northeast-trending faults, consistent with evidence of regional dextral shear along the northern margin of the Rheic Ocean in the Middle Devonian. Late Devonian–early Carboniferous deposition of the predominantly continental clastic rocks of the Horton Group occurred around the Antigonish Highlands in a series of grabens and half-grabens, most notably represented by the St. Mary’s basin, which originated by dextral shear along the boundary between the Meguma and Avalon terranes. Continued episodes of dextral shear in the late Carboniferous resulted in localized regions of transtension and basin development, and also in episodes of transpression, manifested by intense deformation, thrusting, and S-C fabric development. Taken together, Middle Devonian–late Carboniferous episodes of dextral shear reflect the local accommodation of oblique convergence and eventual collision between Gondwana and Laurussia. Keywords: Arisaig Group, Avalon terrane, Rheic Ocean, Appalachian orogen
INTRODUCTION Ordovician to Carboniferous rocks in the Antigonish Highlands, Nova Scotia, were deposited during a crucial interval in the formation of the Appalachian orogen. The Antigonish Highlands lie within the Avalon terrane (Avalonia), which is the largest suspect terrane in the Canadian Appalachian orogen and occupies much of its southeastern flank (Williams, 1979; Williams and Hatcher, 1983; Keppie, 1985). Correlative rocks occur in New England, Britain, and parts of Bohemia and Iberia (Fig. 1; e.g., Nance and Murphy, 1996; Nance and Thompson, 1996; Linnemann et al., 2000; Murphy et al., 2004a,b,c,d; GutierrezAlonso et al., 2005). Although there were intervals when narrow seaways existed (Landing, 1996, 2005), it is generally accepted that Avalonia was one of several terranes (termed peri-Gondwanan terranes; Nance and Thompson, 1996) that was located either along or proximal to the northern margin of Gondwana from the Late Neoproterozoic to the Late Cambrian or Early Ordovician (Fig. 1; e.g., O’Brien et al., 1983; Keppie, 1985; van Staal et al., 1998; Cocks and Torsvik, 2002). The final separation of Avalonia from Gondwana in the Late Cambrian–Early Ordovician coincides with the opening of the Rheic Ocean with Avalonia defining its northern flank (Fig. 2; Cocks and Fortey, 1990; Cocks and Torsvik, 2002; Stampfli and Borel, 2002). Faunal linkages between Avalonia and Baltica are evident by the Late Ordovician (e.g., Williams et al., 1995; Fortey and Cocks, 2003), and paleomagnetic data indicate the proximity of Avalonia to Laurentia by the Early Silurian (Miller and Kent, 1988; Pickering et al., 1988; Trench and Torsvik, 1992; Potts et al., 1993). Thus, as the Rheic Ocean gradually widened during the Ordovician with the increasing separation of Avalonia from Gondwana, the Iapetus Ocean, which was located to the north of Avalonia, gradually diminished as Avalonia converged on Baltica and Laurentia and had virtually disappeared by the Late Ordovician or Early Silurian (Fig. 2; Miller and Kent, 1988; Van der
Voo, 1988; McKerrow and Scotese, 1990; Soper and Woodcock, 1990; McKerrow et al., 1991; Trench and Torsvik, 1992; Potts et al., 1993; Dalziel et al., 1994; Golonka et al., 1994; MacNiocaill and Smethurst, 1994; MacNiocaill et al., 1997; Hodych and Buchan, 1998; Torsvik and Rehnström, 2003). The nature and timing of the accretion of Avalonia to Laurentia-Baltica are fundamental to the understanding of the evolution of the Appalachian-Caledonide orogen and to the evolution of the Iapetus and Rheic Oceans that lie between Laurentia and Gondwana. This article presents a review of the geological evolution of Ordovician to Carboniferous strata in the Antigonish Highlands and shows that they record key events in the evolution of the Iapetus and Rheic Oceans leading to the Carboniferous amalgamation of Pangea. GENERAL GEOLOGY Neoproterozoic–Early Ordovician The Antigonish Highlands (Fig. 3A) are predominantly underlain by Neoproterozoic (ca. 618–608 Ma) arc-related volcanic and sedimentary rocks of the Georgeville Group and synto late kinematic plutons. These Neoproterozoic rocks are a local expression of the voluminous and regionally extensive ca. 630– 540 Ma Avalonian subduction-related tectonothermal activity that occurred along the northern Gondwanan margin (Fig. 4; Murphy et al., 1999b). They are overlain by two coeval Cambrian–Early Ordovician sequences that are related by facies variations: the Iron Brook Group consists of continental to shallow marine clastic rocks and limestones that contain fauna diagnostic of Avalonia (e.g., Landing and Murphy, 1991) and indicate a genetic connection to the northern Gondwanan margin (Landing, 1996). The coeval MacDonald Brook Group contains bimodal, rift-related intracontinental volcanic rocks that reflect the local development of a pull-apart basin (Murphy et al., 1985).
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Figure 1. Map of the north Atlantic borderlands in their pre-Mesozoic drift positions showing the distribution of Avalonia and other Neoproterozoic peri-Gondwanan terranes (modified from Strachan and Taylor, 1990; Nance and Murphy, 1994; Keppie et al., 2003). The portion of Avalonia in Atlantic Canada is identified as West Avalonia, the portion in Britain and Ireland as East Avalonia. Only peri-Gondwanan terranes where Cambrian overstep sequences occur are identified (see Theokritoff, 1979; Keppie, 1985). The inset shows a map of the Appalachian orogen in maritime Canada and Maine. AB—Antigonish basin; AH—Antigonish Highlands; MB—Merigomish basin; MFZ—Minas fault zone (defines the boundary between the West Avalonia and Meguma terranes); SB—Stellarton basin.
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Figure 2. Early Ordovician to Late Silurian paleogeographic reconstructions of the circum-Atlantic region (based on Cocks and Torsvik, 2002; Fortey and Cocks, 2003). The location of Avalonia at 480 Ma is shown, as is the location of the Dunn Point volcanic rocks at 460 Ma.
Geological evolution of middle to late Paleozoic rocks in the Avalon terrane
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Cambrian–Early Ordovician rocks are unconformably overlain by ~80 m of bimodal volcanic rocks and interbedded red clastic sediments variously known as the Dunn Point Formation (north of the Hollow fault; see Fig. 3B) and the correlative Bears Brook Formation (south of the Hollow Fault; see Fig. 3A). The basalts show geochemical characteristics typical of differentiated continental tholeiites, and the felsic rocks were probably generated by crustal anatexis, suggesting that the volcanism was due to local crustal extension (Keppie et al., 1979; Murphy, 1987). Until recently, the only available geochronological datum was an Rb-Sr whole-rock isochron that yielded an age of 421 ± 15 Ma (Fullagar and Bottino, 1968, recalculated in Keppie and Smith, 1978). Although imprecise, these data, together with the apparent lack of significant deformation and lack of discordance with the overlying Arisaig Group strata suggested that the volcanic rocks were a manifestation of the rifting that heralded the development of a basin into which Arisaig Group clastic rocks were deposited (e.g., Murphy, 1987; Waldron et al., 1996). However, recent U-Pb zircon data (Hamilton and Murphy, 2004) from a flow-banded rhyolite in the Dunn Point Formation yield an age of 460.0 ± 3.4 Ma and imply a significant time gap (ca. 17 million years) between extrusion of the volcanics and the onset of deposition of the Early Silurian–Early Devonian Arisaig Group clastic rocks (Fig. 4).
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Figure 3. (A) Summary of the geology of the Antigonish Highlands. For details, see Murphy et al. (1991). (B) Summary geological map of the Arisaig area (modified from Boucot et al., 1974). Location shown in the inset in Figure 3A.
The Arisaig Group (see Boucot et al., 1974, and Murphy, 1987, for details) consists of an 1800–1900 m continuous stratigraphic sequence dominated by shallow marine fossiliferous siliciclastic rocks that unconformably overlies the Dunn Point and Bears Brook Formations and contains Llandoverian to Lochkovian fossils (Figs. 3B and 4). The basal Beechill Cove Formation contains Early Llandoverian fossils and consists of conglomerate, siltstone, and shale deposited in a near-shore environment overlain by interbedded mudstone, shale, and sandstone (Pickerill and Hurst, 1983). The contact with the overlying Middle to Late Llandoverian Ross Brook Formation is gradational. The Ross Brook Formation consists of interbedded graptolite-bearing black shale, siltstone, arenaceous limestone, and ash beds (Boucot et al., 1974; Hurst and Pickerill, 1986; Bergström et al., 1997) and is conformably overlain by Wenlockian interbedded green fissile shale, siltstone, and sandstone and by minor ironstone of the French River Formation, followed by a sequence of laminated shales, minor siltstone, and arenaceous limestone of the Late Wenlockian Doctors Brook Formation. The overlying MacAdam Brook Formation contains Early Ludlovian fossils and consists of interbedded fissile shale, calcareous siltstone, and fossiliferous limestone and minor ash beds and is followed by the Late Ludlovian Moydart Formation, which consists of green interbedded siltstone, shale, mudstone, and limestone (lower member) overlain by subaerial red
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Murphy ca. 300 Ma CLASTIC ROCKS
BASIN
MABOU GROUP
CLASTIC ROCKS LIMESTONE, EVAPORITE
ca. 340 Ma ca. 340 Ma
OVERSTEP SEQUENCE
WINDSOR GROUP HORTON GROUP
CLASTIC ROCKS BIMODAL VOLCANIC ROCKS
ca. 360 Ma
McARRAS BROOK FORMATION
ca. 380 Ma RED BEDS
SYN- TO POSTCOLLISIONAL RIFT
SHALE
ARISAIG GROUP
SANDSTONE ca. 440 Ma
ARC-RIFT
BIMODAL VOLCANIC ROCKS
ca. 460 Ma ca. 500 Ma SHALE
PLATFORM ca. 540 Ma
DUNN POINT FORMATION IRON BROOK GROUP
SANDSTONE McDONALD'S BIMODAL LST VOLCANIC ROCKS BROOK GROUP QUARTZITE
ca. 540 Ma
ARC TO RIFT TRANSITION
RED BEDS GRANITE
ca. 590 Ma
BIMODAL GRANITE VOLCANIC ROCKS
ca. 590 Ma GRANITE
VOLCANIC ARC BASINS
CONTINENTAL MAGMATIC ARC
VOLCANOGENIC TURBIDITES
GEORGEVILLE GROUP CALC-ALKALINE VOLCANIC GRANITOID ROCKS APPINITE
ca. 620 Ma 1.0 Ga
JUVENILE ARC BASEMENT
PROTO-AVALONIA Sm-Nd isotopes
1.2 Ga (not to scale)
Figure 4. Tectonostratigraphy of the Antigonish Highlands and surrounding areas (modified from Murphy et al., 1991; Murphy and Nance, 2002). Names of stratigraphic units exposed in and around the Antigonish Highlands are given at right.
Geological evolution of middle to late Paleozoic rocks in the Avalon terrane mudstones with caliche (upper member, Dineley, 1963; Lane and Jensen, 1975; Waldron et al., 1996). These strata are conformably overlain by the Pridolian Stonehouse Formation, which consists of interbedded mudstones, shales with minor siltstones, and sandstones. The overlying Knoydart Formation consists of interbedded red and green coarse- to fine-grained clastic rocks deposited in deltaic and fluviatile environments (Boucot et al., 1974). Middle–Late Devonian Arisaig Group strata are overlain with angular unconformity by the Middle Devonian McArras Brook Formation (Fig. 3B), which consists of interbedded basalts and continental redbeds in the type area in the northwestern highlands and includes interbedded rhyolites in the northeastern highlands (Boucot et al., 1974). The McArras Brook Formation is only moderately tilted in the type area and displays no internal deformation. In contrast, the structure of the underlying Arisaig Group and Dunn Point Formation is dominated by a large-scale (km), NNE-trending, gently southwest-plunging syncline (Fig. 3B) and a locally penetrative axial planar fabric reflecting deformation prior to the deposition of the McArras Brook Formation (Boucot et al., 1974) and is widely believed to represent a local expression of the Acadian orogeny (Murphy and Keppie, 1998; Braid and Murphy, 2006). Sedimentary strata in the McArras Brook Formation have yielded Late Devonian (Famennian) fossils (Keppie et al., 1978; Martel et al., 1993) and U-Pb (zircon) data on the rhyolites indicate an age of 370+3/–2 Ma (Dunning et al., 2002). The McArras Brook Formation and correlative rocks in the Cobequid Highlands to the west and on Cape Breton Island to the east are all attributed to localized intracontinental rifting accompanied by dextral strike-slip tectonics (e.g., Keppie et al., 1978; Murphy and Keppie, 1998; Piper and Pe-Piper, 2001). Thick deposits of Upper Devonian to late Carboniferous strata around the Antigonish Highlands unconformably overlie Neoproterozoic to Middle Devonian rocks. These strata are deposited along the southern flank of the composite Maritimes basin, which contains several depocenters, and are subdivided into three allocycles (e.g., Ryan et al., 1987; Gibling et al., 1992): (1) the uppermost Devonian—Tournaisian, predominantly nonmarine, siliciclastic Horton Group, overlain by (2) the Visean, predominantly marine, Windsor Group, followed by (3) Namurian–Westphalian clastic sequences that include coal deposits (the Mabou Group). Upper Devonian–Tournaisian The Horton Group, which unconformably overlies the McArras Brook Formation, is widespread throughout maritime Canada and appears to have been deposited in a series of grabens or half-grabens (McCutcheon and Robinson, 1987; Hamblin and Rust, 1989; Martel and Gibling, 1996). The Horton Group is dominated by up to 4000 m of continental clastic deposits ranging from conglomerates to shales that contain latest Devonian–
419
Tournaisian plant fossils and spores (e.g., Utting and Hamblin, 1991; Martel et al., 1993). In the northern and eastern Antigonish Highlands, the Horton Group contains abundant clasts that are clearly derived locally from the underlying Antigonish Highlands rocks. However, they also contain ca. 370 Ma detrital muscovites attributed to erosion of the Meguma terrane to the south (Dallmeyer et al., 1997), which, unlike the Avalon terrane, has abundant muscovite-bearing rocks of that age. Taken together, these data imply a mixed Avalon-Meguma provenance. Along the southern flank of the Antigonish Highlands, Horton Group rocks in the St. Mary’s basin (SMB in Fig. 1) consist of an intracontinental alluvial fan-fluviatile-lacustrine basin-fill sequence that occurs within the east-west-trending Minas fault zone (MFZ in Fig. 1) that forms the boundary between the Avalon and Meguma terranes. The evolution of this basin reveals much about Avalon-Meguma relationships in the late Paleozoic (Fig. 5; Murphy, 2003). Field observations as well as geochemical and U-Pb isotopic data on detrital zircons clearly indicate that the vast majority of the detritus was derived from the Meguma terrane (Murphy et al., 1995b; Jennex et al., 2000; Murphy, 2000; Murphy and Hamilton, 2000; Murphy, 2003). A strong tectonic influence on sedimentation is apparent along the entire southern (Meguma) flank of the basin, where subsidence along northerly-dipping listric normal faults is associated with deposition of coarse conglomerates. In contrast, the character of the sediments does not vary with proximity to the northern margin (the Chedabucto fault), suggesting that the fault does not constitute the original basin margin and that an unknown portion of the basin has been tectonically removed (Murphy et al., 1995b). Deformation of St. Mary’s basin rocks is very heterogeneous, and the most intense deformation within the basin is concentrated in a relatively narrow ENE-trending zone, in which predominantly fine-grained clastic rocks are deformed into periclinal folds and related reverse faults (Fig. 5). The orientation of this zone relative to the Minas fault zone is consistent with dextral shear (Murphy et al., 1995b; Murphy, 2003). An angular unconformable relationship between the Horton Group and the overlying Windsor Group is indicated by a gap in the depositional record across the contact in the western portion of the St. Mary’s basin (Murphy, 2003). However, northeast-oriented folds can be mapped across this contact, suggesting that at least some of the deformation was post-Tournaisian in age (Murphy, 2003). This interpretation is compatible with 40 Ar-39Ar (muscovite) data from metasedimentary rocks within the Meguma terrane that yield ages ranging from 370 to 300 Ma, suggesting that episodes of deformation occurred at various times in the Carboniferous (e.g., Keppie and Dallmeyer, 1995; Culshaw and Liesa, 1997). Visean–Namurian Visean rocks are represented by the Windsor Group, which typically consists of 750–1000 m of cyclical alternations of
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Murphy 63 00' R oc k land Br o ok F a ul t Por tapique Faul t
Minas Basin
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Mesozoic rift sediments and basalt Windsor Group Liscomb Complex (Meguma) Precambrian (Avalonia)
45 25'
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M E G U M A
A
Erosion Level Stratigraphic contact Fault Fold, with plunge
Figure 5. (A) Summary geological map of St. Mary’s basin (after Murphy, 2003) showing the distribution of Horton Group clastic rocks and regional fold hinges. Note the ENE-trending hinge lines, attributed to dextral shear along the bounding faults of the basin. Stratigraphy: B—Barrens Hills; C—Cross Brook; CHF—Country Harbour fault; GH—Graham Hill; L—Lochiel; LS—Little Stewiacke River; WR—West River St. Mary’s. For location, see Figures 1 and 3A. The inset shows typical facies (modified from Murphy and Rice, 1998). (B) Deformation in a Meguma granitoid rock along the contact of the basin with the Meguma terrane. C-S fabrics in granite show evidence of dextral shear. (C) Moderately plunging stretching lineations in the same outcrop as pictured in B.
limestones or dolostones; evaporites, including gypsum, anhydrite, salt, and potash; and fine- to coarse-grained clastic rocks (Figure 4; e.g., Boehner and Giles, 1993). The base of the Windsor Group appears to mark a very rapid, basinwide transgressive event and the cyclical nature of the sedimentation suggests repeated transgressive-regressive events in a predominantly arid climatic setting (Giles, 1981; Ryan and Boehner, 1994; Chandler, 1997). Although regionally extensive, the Windsor Group is deposited in several local depocenters that are components of the composite Maritimes basin (e.g., Boehner and Giles, 1993). Two subbasins occur around the Antigonish Highlands, the Merigomish sub-basin to the northwest and the Antigonish sub-basin to the east (Fig. 1). In the Merigomish sub-basin, the Windsor Group is
divided into continental clastic rocks of the Martin Road Formation overlain by laminated limestone of the Ardness Formation (Fig. 6). The Martin Road Formation lies unconformably on the McArras Brook Formation and consists of ~120 m of upwardfining redbeds (Keppie et al., 1978) that contain Visean AT plant fossils (Chandler, 1997). Geochemical and isotopic analyses indicate that the Martin Road Formation is primarily locally derived from the underlying Arisaig Group strata (Murphy, 2001). The Ardness limestone is disconformably overlain by the Namurian Lismore Formation, which is a local representative of the regionally extensive Mabou Group. The Ardness limestone contains Visean C to E1 spores, suggesting that it is one of the youngest limestone horizons within the Windsor Group. Spore
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Chedabuc to
Stellarton Basin
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Jessie's Cove
McARRAS BROOK FM
BASEMENT ROCKS OF THE ANTIGONISH AND COBEQUID HIGHLANDS
DEVONIAN & EARLIER ROCKS
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FOUNTAIN LAKE GROUP
LATE DEVONIAN ROCKS
FALLS FORMATION
DEVONO-CARBONIFEROUS ROCKS
NUTTBY FORMATION
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ARDNESS FORMATION & MARTIN ROAD FORMATION
WINDSOR GROUP UNDIVIDED
MIDDLEBOROUGH FORMATION
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MABOU GROUP
LISMORE FORMATION
PARRSBORO FORMATION
CLAREMONT FORMATION
NAMURIAN ROCKS
BOSS POINT FORMATION
MIDDLE RIVER FORMATION
NEW GLASGOW FORMATION, conglomerate member NEW GLASGOW FORMATION, lower member
UNDIVIDED WESTPHALIAN ROCKS
STELLARTON FORMATION
MALAGASH FORMATION
PICTOU GROUP
WESTPHALIAN-STEPHANIAN ROCKS
LEGEND
Figure 6. Summary geological map (after Chandler et al., 1997; Stevens et al., 1999) of the Merigomish subbasin study area within the Maritimes basin (inset) between the Antigonish Highlands (AH) and the Cobequid Highlands (CBH). CH—Caledonia Highlands; DH—drillhole location; FM—Formation; MT—Meguma terrane. N, NB, VAT, VSM, and WC are spores of Visean (B), Namurian (N), and Westphalian (W) ages.
C obequid Highlands
Sc o
t aul nF
Meadowville
Toney River
Pictou Island
Geological evolution of middle to late Paleozoic rocks in the Avalon terrane 421
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Figure 7. Plots showing mixing curves between the late Devonian granitic rocks of the Cobequid Highlands (COBP), the sedimentary rocks of the Arisaig Group (AG), the Meguma Group metasedimentary rocks (MMS), and the distribution of the Lismore Formation sedimentary rocks. (A) Zr/Y versus Ti/Nb, (B) Zr/Y versus Ti/V, and (C) V/Nb versus Zr/Y (all in parts per million). Modified from Stevens et al. (1999).
V/Nb
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COBP
.1
C .01 1
10 Zr/Y
100
samples from Lismore Formation redbeds directly overlying the limestone suggest a Visean AT age (upper Visean; Utting, 1987), whereas farther upsection, the Lismore Formation has yielded early Namurian and Namurian B spore assemblages (Chandler et al., 1997). The Lismore Formation consists of 2500 m of upwardcoarsening fluvial strata and is divided into upper and lower members that reflect variations in depositional environment and paleoclimate. Geochemical and isotopic data (Stevens et al., 1999) indicate that the Lismore Formation can be subdivided into two groupings that primarily reflect varying contributions from different local source rocks, as reflected by accessory phases, clay minerals, or rock fragments (Fig. 7). This subdivision occurs 150 m above the base of the upper member. The data from the lower grouping (group A) show a predominant contribution from underlying Silurian Arisaig Group rocks, whereas the data from the upper grouping (group B) reveal an additional contribution from ca. 360 Ma Cobequid Highlands granitoid rocks (see location of Cobequid Highlands in Fig. 1). This change in geochemistry is attributed to renewed motion and uplift along the faults along the southern flank of the Maritimes basin during Lismore Formation deposition (Fig. 7). Support for such uplift is derived from kinematic studies of major
fault zones, such as the Rockland Brook fault zone in the Cobequid Highlands, which underwent dextral transpression during that time (Miller et al., 1995). In the Antigonish sub-basin, the Windsor Group consists of a basal mid-Visean finely laminated limestone and horizons of brecciated limestone (the Macumber Formation), 3 to 25 m in thickness, that rims the basin margin (Fig. 8A). These are overlain by a predominantly evaporitic sequence of halite gypsum and anhydrite, with minor limestone and siltstone followed by interbedded limestones, anhydrites, siltstones, and shales that have been interpreted as recording a series of transgressive-regressive events (Boehner and Giles, 1993). The limestone beds are highly deformed, with inclined to recumbent folds that do not affect the underlying finely laminated horizon at the base of the formation or the underlying Horton Group (Fig. 8A–C). These folds show broadly parallel geometries, are easterly vergent, and have amplitudes that increase from 2 to 3 cm near the base of the deformed zone to 40 to 50 cm upsection. The detached style of the folding suggests an origin through slumping, which is supported by the presence of previously folded fragments in overlying breccia units (Fig. 8C). These relationships suggest that at least some of the deformation was syndepositional with respect to the Windsor Group (Thomas et al., 2002).
Geological evolution of middle to late Paleozoic rocks in the Avalon terrane
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St. George’s Bay Monk’s Head Fault
Morristown Fault
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CUMBERLAND GRP.
Mabou Group - Late Carboniferous Upper Windsor Group Lower Windsor Group (Macumber Formation) Devonian - Early Carboniferous Neoproterozoic Georgeville Group
Figure 8. (A) Simplified geological map of the Antigonish basin (modified from Boehner and Giles, 1982). See Figure 3A for location. (B) Schematic representation of stratigraphic relationships within the Antigonish basin (after Boehner and Giles, 1982). (C) Large, east-verging recumbent fold in Macumber Formation breccia at MacIsaac Point. (D) Folded layering within individual clast of limestone breccia in the same formation. C and D after Thomas et al., 2002.
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Westphalian-Stephanian To the west of the Antigonish Highlands, correlatives of the Lismore Formation (the Hollow Conglomerate) were deformed and unconformably overlain by Westphalian strata in the Stellarton basin (see location in Figs. 1 and 6). The evolution of this basin has been outlined in detail by Waldron (2004), and it consists of ~3000 m of late Carboniferous predominantly clastic rocks, coal seams, and oil shales deposited in lacustrine and deltaic settings. Basin development is attributed to local transtension at the stepover between the Cobequid and Hollow faults during regionally extensive dextral strike-slip tectonics (Yeo and Ruixiang, 1987; Waldron, 2004). Implicit in this interpretation is that the Antigonish and Cobequid Highlands were one lithotectonic block prior to the opening of the Stellarton basin, an interpretation supported by correlations between their respective Neoproterozoic strata (e.g., Murphy et al., 1991). Abundant soft-sediment deformation structures are interpreted by Waldron (2004) to reflect coeval tectonic activity. East- to northeast-trending folds are attributed to dextral strike-slip motion on adjacent faults, which also resulted in rotational progressive strain, followed by the development of a positive flower structure along the northwest margin, representing a transition from transtension to transpression (Waldron, 2004). To the northeast of the Antigonish Highlands near the Cape George Peninsula, unmetamorphosed Paleozoic sequences dominated by Arisaig Group, McArras Brook Formation, and Horton
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Figure 9. Compilation (based on MacNiocaill, 2000) showing the predicted paleolatitude of Avalonia and the Newfoundland margin of Laurentia during the development of the Rheic Ocean. West Avalonia: DP— Dunn Point; CSM—Cape St. Mary’s; SW—Springdale and Wigwam. East Avalonia: TrV—Treffgarne volcanics; SV—Stately volcanics; BV—Builth volcanics; NB—North Builth volcanics/sediments/intrusives; TV—Tramore volcanics; BoV—Borrowdale volcanics; TL— Tortworth volcanics; BR—Browgill volcanics; MV—Mendip volcanics; MC—Mill Cove redbeds; OR—Old Red Sandstone. For a compilation of references, see MacNiocaill (2000). For discussion on Dunn Point, see Hamilton and Murphy (2004); Johnson and Van der Voo (1990).
Group strata display anomalously intense brittle to ductile deformation, including thrusts, and local development of S-C fabrics and stretching lineations. The orientation of these structures is consistent with dextral transpression along the adjacent Hollow fault, and their age is thought to be virtually coeval with the origin and development of the Stellarton basin (St. Jean et al., 1993). REGIONAL SETTING The Antigonish Highlands are part of Avalonia or the Avalon terrane (Keppie, 1985; Landing and Murphy, 1991; Murphy et al., 1991; Landing, 1996). Paleomagnetic and faunal evidence (e.g., Johnson and Van der Voo, 1986; Landing, 1996; McNamara et al., 2001) and tectonic syntheses (e.g., O’Brien et al., 1983; Keppie, 1985; Murphy and Nance, 1989, 1991) indicate that Avalonia was geodynamically linked to the northern margin of Gondwana from the Neoproterozoic to the Late Cambrian, although seaways may have existed that allowed faunal endemism (Landing, 1996, 2005). Reviews of available paleomagnetic and faunal data (Cocks and Torsvik, 2002; Fortey and Cocks, 2003) indicate that Avalonia had separated from Gondwana by the Early Ordovician to form a microcontinent that migrated from high latitude (~60°S) to low latitude (20°S) during the Ordovician (Fig. 2) as the Rheic Ocean to its south widened and the Iapetus Ocean to the north narrowed (Fig. 9; MacNiocaill and Smethurst, 1994). By the Llandeilo, Avalonian brachiopods and trilobites began to show increasing similarity to those of Baltica and North America, suggesting that Avalonia migrated from the cold latitudes of the Gondwanan margin toward more warm temperate paleolatitudes (Fortey and Cocks, 2003). Although proximity to Baltica is indicated by similar fauna in the Caradoc, Avalonia retains some endemic fauna and some that have affinities with Gondwana. Taken together, these data are generally interpreted to reflect Avalonia’s microcontinental position between Gondwana, Laurentia, and Baltica in the early Late Ordovician (Lees et al., 2002). At the end of the Ordovician, Avalonia docked with Baltica, and in the Middle Silurian with Laurentia, where it lay in an outboard position along Laurussia (Murphy et al., 1995a; Cocks and Torsvik, 2002; Fortey and Cocks, 2003). Faunal distinctions between Avalonia and Laurentia-Baltica in the Ashgill were relatively minor (Williams et al., 2001), including increasing exchanges of ostracods. By the Llandovery, Avalonia shared the same brachiopod species with Laurentia or Baltica (e.g., Cocks and Worsley, 1993). Interpretation of the Siluro-Devonian history of Avalonia depends on the understanding of its relationship with the Meguma terrane. Many workers (e.g., Schenk, 1997; van Staal et al., 1998) interpret from the different Cambrian–Early Ordovician lithostratigraphies that the Meguma and Avalon traveled from Gondwana as separate terranes and accreted at different times to the Laurentian margin. In this scenario, the accretion of the Meguma terrane occurred after that of Avalonia, an event commonly held to be responsible for the Middle Devonian Acadian orogeny. In contrast, Keppie et al. (1997),
Geological evolution of middle to late Paleozoic rocks in the Avalon terrane EASTERN LAURENTIA
BALTICA
AVALONIA NS NB NE
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Orogenesis Plutons Rifting Volcanics Sediments
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Figure 10. Comparison of detrital zircon ages from Bears Brook Formation and Arisaig Group rocks with those from the Annapolis Group of the Meguma terrane (Krogh and Keppie, 1990), and Avalonia (Nova Scotia, Keppie et al., 1998; New Brunswick, Bevier et al., 1990; New England, Karabinos and Gromet, 1993). The x’s indicate concordant U-Pb zircon ages; the filled circles (•) indicate discordant 207Pb/206Pb ages; the v’s indicate volcanic rocks. Also shown are tectonothermal events in Baltica (Gower et al., 1990; Goodwin, 1991; Starmer, 1993; Roberts and Siedlecka, 2002; Roberts, 2003), eastern Laurentia (Hoffman, 1989; Connelly and Heaman, 1993; McLelland et al., 1993; Gower and Tucker, 1994; Martignole et al., 1994; Cawood et al., 2001), the Amazon craton (Teixiera et al., 1989; Sadowski and Bettencourt, 1996, and references therein), northwest Africa (Rocci et al., 1991), the Meguma terrane (Krogh and Keppie, 1990), and the Gander Terrane (van Staal et al., 1996, 1998).
Murphy et al. (2004a,c), and Murphy and Keppie (2005) point out that the Sm-Nd isotopic signatures of Early Silurian crustally derived felsic volcanic rocks of the White Rock Formation in the Meguma terrane have a distinctive Avalonian signature, implying that Avalonian crust resided beneath the Meguma terrane by the Early Silurian. In addition, coeval clastic rocks contain an important Mesoproterozoic zircon population that is not present in the underlying Meguma Group, implying that the Meguma terrane could not have been a microcontinent at that time (Fig. 10). Given that Avalonian strata contain abundant Mesoproterozoic zircons (e.g., Keppie et al., 1998; Murphy et al., 2004b), the Sm-Nd and U-Pb (zircon) data imply that the Meguma and Avalon terranes shared a common history by the start of the Silurian. However, because the Meguma Group
contains no record of pre-Silurian penetrative deformation (the oldest deformation recorded is ca. 405–415 Ma; Keppie and Dallmeyer, 1995; Hicks et al., 1999), this common history may have extended as far back as the Neoproterozoic. This interpretation of a contiguous relationship between Avalonia and Meguma implies that they accreted to the Laurentian margin together (by the Llandovery). Because the Meguma is the most outboard terrane in the Appalachians, this interpretation implies that the Middle Devonian Acadian orogeny could not have been a collisional event; instead it is attributed to Andean-style subduction (Murphy et al., 1999a, 2004a; Murphy and Keppie, 2005). In this context, the Middle to Late Devonian deformation is attributed to the effects of dextral strike-slip along the Minas fault zone that now separates the two terranes.
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The post-Acadian lithostratigraphy of Avalonia involves widespread deposition of Late Devonian–late Carboniferous rocks that can be correlated across pre-existing terrane boundaries in Maritime Canada (Williams, 1979; Keppie, 1985). The Tournaisian Horton Group predominantly consists of intracontinental clastic strata deposited in a number of grabens or half-grabens (e.g., Hamblin and Rust, 1989). Although graben development clearly represents lithospheric extension, the mechanism for graben formation is unclear. A widespread marine incursion during the Visean led to the development of a shallow epicontinental sea, as recorded by the deposition of the limestones and evaporate sequences (salt, gypsum, anhydrite, potash) of the Windsor Group (Boehner and Giles, 1993; Boehner, 1994). The late Carboniferous is characterized by the deposition of intracontinental clastic deposits (including deposition of coal seams) in a large depocenter, a key stage in the development of the Maritimes basin (Fig. 11). Gibling et al.
(1992) interpret the paleocurrent data to reflect mountain-building processes that were occurring to the west associated with the Alleghanian orogeny and the amalgamation of Pangea. SYNTHESIS All the key events in the evolution of the Iapetus and Rheic Oceans leading to the Carboniferous amalgamation of Pangea are recorded in the middle to late Paleozoic strata in the Antigonish Highlands. Avalonia as a Microcontinent Paleomagnetic data from Dunn Point mafic rocks (Van der Voo and Johnson, 1985; Johnson and Van der Voo, 1990) imply that Avalonia was located ~10° north of Gondwana and 20° south of Laurentia at 460 Ma, implying that Avalonia was a
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Figure 11. Map of the Maritimes basin showing the distribution of Late Devonian–Tournaisian Horton Group and late Carboniferous– Permian clastic sediments in various depocenters in Atlantic Canada (modified from Martel, 1987; St. Peter, 1993) and a cross-section (A–A′) derived from reflection seismic data (line 86-1 of Marillier et al., 1989). HF—Hollow fault; MFZ—Minas fault zone; MM— Moncton-Malpeque Basin; SB—Sydney basin; SMSB—St. Mary’s Basin.
Geological evolution of middle to late Paleozoic rocks in the Avalon terrane
Collision with Baltica Other than faunal similarities, the best evidence for the timing of proximity to Baltica is derived from the geochemical and Sm-Nd isotopic signature of the Arisaig Group sedimentary rocks, especially the lowermost Beechill Cove Formation, which are so distinct from the underlying Avalonian basement in major, trace, and rare earth elemental abundances that a significant chemical contribution from the Avalonian basement rocks is precluded (Murphy et al., 1996a, 2004c). These distinctions are most clearly evident, however, in the contrasting Sm-Nd isotopic signatures between Arisaig Group sedimentary rocks and underlying strata. Arisaig Group sedimentary rocks are characterized by strongly negative εNd (Fig. 12A and B) and TDM ages of ca. 1.5 Ga, i.e., more than 1.0 Ga older than the depositional age. The Beechill Cove Formation has εNd ranges from −4.8 to −6.1,
and there is an overall trend toward increasingly negative εNd values from the base to the top of the group, with εNd ranging from −6.0 to −7.9 in the Ross Brook, French River, and Doctors Brook formations and from −7.6 to −9.3 in the Stonehouse Formation (Murphy et al., 2004c). These data contrast with those from the immediately underlying Dunn Point volcanic rocks and the Late Neoproterozoic (Georgeville Group) sedimentary rocks, which also contain higher εNd (ranging from +0.16 to +4.39, calculated for a 615 Ma depositional age) and younger εNd (T ) model ages (0.96–1.15 Ga) that are DM typical of Avalonian crust (Murphy and MacDonald, 1993; Mur-
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microcontinent that moved northward at ~8 cm/yr between 485 and 460 Ma (Hamilton and Murphy, 2004). This Arenig-Llandeilo northward drift is consistent with the presence of an ocean ridge system between Avalonia and Gondwana, implying that the trailing (southern) edge of Avalonia would represent a passive margin. However, coeval arc-related magmatism is common in other areas of Avalonia and in Iapetan terranes (van Staal et al., 1998; Woodcock and Strachan, 2000), including voluminous Late Tremadoc–Caradoc arc (e.g., Popelogan-Victoria) and back-arc (Tetagouche-Exploits) complexes in Newfoundland, New Brunswick, and Maine and the ensialic arc magmatism recorded by similar-aged Avalonian rocks in Britain and Ireland (the Lake District–Leinster volcanic arc). These data are consistent with models requiring subduction of Iapetan oceanic lithosphere beneath the northern margin of Avalonia at 460 Ma (e.g., Bevins et al., 1992; van Staal et al., 1998; Woodcock and Strachan, 2000; Torsvik and Rehnström, 2003). In this context, the Dunn Point volcanism may reflect a rift within a continental arc setting that was outboard from both Gondwana and Laurentia, possibly analogous to the modern Taupo volcanic zone in northern New Zealand, which separated from Australia during the Cretaceous opening of the Tasman Sea. New Zealand lies along an obliquely convergent plate boundary where the Pacific plate is being subducted westward under the Australia plate at a rate of ~5 cm/year to form the Taupo-Hikurangi arc-trench system (Cole et al., 1995, 2001). Dunn Point felsic volcanic rocks have Avalonian signatures, with high εNd (ranging from −0.10 to +5.07 at 430 Ma age) and low Nd (depleted mantle, or T ) model ages (0.88–1.05 Ga), DM consistent with the interpretation of these rocks as crustal melts (Murphy et al., 1996b). The detrital zircon populations in the continental red arkosic sandstones of the Bears Brook Formation (which is coeval with the Dunn Point Formation) are remarkably similar to those in the underlying Cambrian rocks (Murphy et al., 2004b,c) in the Antigonish Highlands and are overwhelmingly dominated by Neoproterozoic zircons (56 in a total of 62) between ca. 580 Ma and 680 Ma. These data are consistent with the interpretation of Avalonia as a microcontinent at 460 Ma.
427
S EW
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t Figure 12. (A) εNd versus 147Sm/144Nd diagram (t = 430 Ma) comparing Sm-Nd isotopic data for the Arisaig Group with typical Sm-Nd isotopic compositions of Avalonian crust (Murphy and MacDonald, 1993; Murphy et al., 1996a). The Sm-Nd isotopic characteristics for the average upper crust are bracketed between modern global average river sediment (147Sm/144Nd = 0.114; depleted mantle [TDM] = 1.52 Ga; Goldstein and Jacobsen, 1988) and the average age of sedimentary mass (Miller et al., 1986). See also Thorogood (1990). Iapetan crust includes normal and depleted island arc tholeiites and ophiolitic complexes in Newfoundland and Norway (Pedersen and Dunning, 1997; MacLachlan and Dunning, 1998). Silurian strata of England and Wales (SEW) from Thorogood (1990) and Meguma terrane metasedimentary rocks t (MMS) from Clarke et al. (1997). (B) εNd versus time (Ga) diagram (t = 430 Ma) comparing Sm-Nd isotopic data for the Arisaig Group with typical Sm-Nd isotopic compositions of Avalonian crust (Murphy et al., 1996a; Murphy et al., 2000). CHUR—chondrite uniform reservoir. Field for Grenville rocks after Dickin and McNutt (1989); Patchett and Ruiz (1989); Dickin et al. (1990); Daly and McLelland (1991); McLelland et al. (1993); Dickin (2000).
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phy et al., 1996a, Murphy and Nance, 2002). The contrasting Sm-Nd isotopic signatures indicate that the Arisaig Group clastic rocks were derived either directly or indirectly from basement sources more ancient than Avalonian basement. The clastic rocks of the Arisaig Group contain a very wide range of detrital zircon populations, including some Ordovician-Silurian zircons that are similar to the depositional age of the respective strata (Murphy et al., 2004a,c), suggesting derivation from coeval volcanic rocks. Silurian volcanism is common in the Appalachians (e.g., Barr et al., 2002), including in terranes inboard of Avalonia that had previously accreted to Laurentia (e.g., van Staal et al., 1998). Llandovery and Ludlovian K-bentonite ash beds in the Arisaig Group are thought to be related to the regionally extensive volcanism (Bergström et al., 1997). The U-Pb detrital zircon ages in the Arisaig Group are consistent with derivation from eastern Laurentia, Baltica, or Amazonia, which are all potential sources of the ca. 2.1–0.9 Ga zircons (Fig. 10). The Meguma terrane is an unlikely source, for it underwent greenschist- to amphibolite-facies metamorphism during Arisaig Group deposition, implying that it was not at the surface during Arisaig Group deposition and so cannot be a source of the Arisaig Group detritus. Ordovician–Silurian continental reconstructions (Fig. 2) suggest that direct derivation from Amazonia is unlikely, and the strongly negative Sm-Nd isotopic signature negates the possibility that significant amounts of Amazonian zircons were recycled from Avalonian sediments into Arisaig detritus. Most reconstructions (e.g., Fig. 2) show Avalonia adjacent to Baltica during deposition of the lowermost formations of the Arisaig Group and adjacent to Laurentia by the time of deposition of the uppermost formations. Baltica could have provided the bulk of the zircons deposited in the Beechill Cove sample, an interpretation compatible with the paleocurrent data in the Arisaig Group (Boucot et al., 1974), which show derivation from the northeast (present coordinates). If so, the lower part of the Arisaig Group is probably part of a sequence that overstepped the boundary between Avalonia and Baltica. The lack of significant deformation in the underlying Dunn Point volcanics and the minor angular discordance between these volcanics and the Beechill Cove Formation are consistent with the oblique “soft” collision between Avalonia and Baltica, in which the Arisaig Group was deposited near the trailing edge of the Avalonia plate. Collision with Laurentia By the time of deposition of the ca. 425 Ma French River and ca. 400 Ma Stonehouse Formation in the early Devonian, Avalonia-Baltica had collided with the eastern Laurentian margin, resulting in deformation and uplift of the Laurentia-Baltica margin and of terranes inboard of Avalonia (e.g., the Gander terrane; van Staal et al., 1990, 1996, 1998). Such deformation would account for the Late Silurian (ca. 415 Ma) loading of the Avalonian margin documented by Waldron et al. (1996), which
reflects the development of a foreland basin. Alternatively, loading and foreland basin development could be attributed to dextral transpression of the Meguma terrane relative to the Avalon terrane along the Minas fault zone. If the Meguma and Avalon terranes had been contiguous from the start of the Paleozoic (Murphy et al., 1999a, 2004a; Murphy and Keppie, 2005), this would reflect intracontinental dextral shear analogous to the retro-arc region of the Andes. If the Meguma and Avalon terranes had separate early Paleozoic histories (van Staal et al., 1998), loading might reflect the collision of the Meguma terrane with the Avalonian margin of Laurentia. The more negative εNd values of the Stonehouse Formation suggest that the loading of Avalonia was coeval with the exposure of older basement and are more consistent with derivation from terranes inboard of Laurentia and the Laurentian margin rather than the Meguma terrane, which was undergoing metamorphism at the time of Stonehouse Formation deposition. In addition, because the Meguma Group is dominated by ca. 600 Ma and 2.1 Ga zircon populations (Krogh and Keppie, 1990), derivation from the Meguma Group cannot account for the zircon populations in the Stonehouse Formation. Acadian orogeny Although the deepening of the Arisaig basin heralded the onset of the Acadian orogeny, the Arisaig Group was not deformed until the Middle Devonian, when the Arisaig Group strata formed regional northeast- to NNE-trending folds, and recorded deformation processes in the relatively shallow crust (Fig. 3B; Boucot et al., 1974). The deformation predated the deposition of the internally undeformed and unconformably overlying Middle Devonian McArras Brook Formation. The observed structural features are consistent with fold propagation associated with ramp–flat thrust fault geometry and coeval local extension recorded by a set of conjugate normal faults (Braid and Murphy, 2006). Many outcrop-scale folds have sheared limbs and show evidence of a complex progressive deformation (Fig. 13). The orientations of the folds and subsequent faults are <45° to the adjacent Hollow fault (Fig. 14), consistent with the orientation of the infinitesimal strain ellipse for simple shear in a dextral strike-slip zone, with normal faults oriented at a high angle to the Hollow fault. These geometric relationships are consistent with a single progressive deformation instigated by dextral motion along the Hollow fault. Irrespective of whether Meguma and Avalonia had contiguous or separate early Paleozoic histories, deformation of the Arisaig Group is considered a far-field effect of dextral translation of the Meguma terrane relative to Avalonia along the Minas fault zone (Braid and Murphy, 2006). Late to Post-Acadian History and the Amalgamation of Pangea There is abundant evidence of repeated episodes of dextral shear from the Late Devonian to the end of the Carboniferous, by
Geological evolution of middle to late Paleozoic rocks in the Avalon terrane
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FOLD
Figure 13. Field photo from the upper Ross Brook Formation showing a typical outcrop-scale fold that is propagated along the leading edge of a thrust (see Braid and Murphy, 2006).
THRUST
which time Pangea had amalgamated. This dextral shear regime is consistent with the clockwise rotation of Laurussia relative to Gondwana during the Carboniferous (e.g., Keppie, 1985; Dalziel, 1997). The St. Mary’s basin (Figs. 3B and 5) is an excellent example of basin development and evolution adjacent to an intracontinental fault zone associated with oblique convergence during orogenesis (Murphy, 2003). Its evolution provides constraints on the potential relationship between the termination of the Acadian orogeny and subsequent basin development during the ongoing dextral translation along the Avalon-Meguma terrane boundary, and on the relationship between Laurentia and Gondwana during the assembly of Pangea. Between ca. 375 and 360 Ma, rapid uplift of the Meguma terrane relative to the Avalon terrane is recorded in Horton Group strata within the St. Mary’s basin. Kinematic evidence within Meguma terrane lithologies adjacent to the St. Mary’s basin indicates that uplift was accompanied by dextral shear along the Minas fault zone (Keppie and Dallmeyer, 1995; Murphy, 2003). Just south of the basin, granites typical of the ca. 370 Ma Meguma terrane granitoid rocks have well-developed S-C fabrics and subhorizontal stretching lineations that indicate dextral shear (Fig. 5B and C; Murphy, 2003). The lack of deformation in the adjacent Horton Group rocks indicates that this deformation occurred between ca. 370 and 365 Ma, and therefore probably heralds basin development. The abundance of Meguma terrane detritus in St. Mary’s Basin Horton Group rocks is indicated by (1) the presence of clasts of metasediments, two-mica granitoid rocks, vein
quartz, and gold, which are all characteristic of the Meguma terrane (Murphy et al., 1995b; Jennex et al., 2000); (2) lithogeochemical, Sm-Nd isotopic, and U-Pb detrital zircon data (Murphy and Hamilton, 2000; Murphy, 2003) suggesting that dextral shear was accompanied by uplift of the Meguma terrane relative to Avalonia; and (3) 40Ar/39Ar data from Meguma terrane rocks that indicate 5–12 km of uplift relative to Avalonia between 370 and 360 Ma (Keppie and Dallmeyer, 1995). 40 Ar/39Ar data from muscovites in dextral shear zones near the contact with the St. Mary’s basin indicate that cooling through ~400–350 °C continued until ca. 345 Ma (Keppie and Dallmeyer, 1995), suggesting that uplift and regional dextral shear of the Meguma terrane relative to the Avalon terrane continued during St. Mary’s basin deposition. Deformation of the basin-fill rocks is most pronounced in a narrow (2-km wide) zone dominated by less competent fine-grained clastic rocks and produced ENE-trending periclinal folds and associated reverse faults (Murphy, 2003). The orientation of these structures relative to the eastwest-trending Minas fault zone is consistent with their origin by coeval dextral shear along the fault zone (Fig. 5). To the north, the rotation of these structures into parallelism with the Chedabucto fault could be attributed to a progressive continuation of the same event or could reflect a later phase of dextral shear. Horton Group clastic rocks unconformably overlying the Antigonish Highlands are characterized by clasts typical of Antigonish Highland lithologies, including abundant volcanic and low-grade sedimentary clasts that are readily matched with underlying lithologies. Although most of the clasts are locally derived, the rocks also contain abundant muscovite, whose ca. 370 Ma age
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is typical of a Meguma terrane derivation and is atypical of the Avalonia terrane, which was not metamorphosed at that time. The relationship between the Horton Group and the overlying Windsor Group is highly variable within maritime Canada. Regional stratigraphic and paleontological studies indicate a gap in the depositional record across this contact (e.g., Boehner and Giles, 1993; St. Peter, 1993; Boehner, 1994), including in the western portion of the St. Mary’s basin (Murphy, 2003). Contact relationships vary from concordant in some sections to pronounced angular unconformity in others (Fig. 14). These relationships are compatible with the fault-related heterogeneous style of deformation of the Horton Group strata, depending on fault geometries and proximity of active faults. Angular unconformities between the Horton and Windsor Groups are especially pronounced near basin margins, including the eastern flank of the Antigonish Highlands (Boehner and Giles, 1993) and the western margin of the St. Mary’s Basin (Murphy et al., 1995b). These variable relationships suggest a relationship between preWindsor deformation and motion along basin-bounding faults (Fig. 14; Murphy, 2003). The Windsor Group represents a period of widespread flooding. In the Merigomish sub-basin, the presence of only one thin and relatively young limestone deposit indicated progressive onlap of Windsor seas onto the northern margin of the Antigonish Highlands. The origin of the intraformational slump folding of Windsor Group lithologies in the Antigonish basin could be due to salt withdrawal and/or tectonic instability along the edges of the basin during Windsor Group deposition (Thomas et al., 2002). Salt bodies of presumed Windsor Group age have been interpreted in seismic sections in the central portions of the basin that underlie St. George’s Bay (Durling et al., 1995a,b), whereas they are not identified in on-land exposures at the basin margins. By the late Carboniferous, most reconstructions (e.g., McKerrow and Scotese, 1990; Golonka et al., 1994; Stampfli and Borel, 2002) suggest that Gondwana and Laurussia had collided to form Pangea, implying that maritime Canada lay near the center of a large supercontinent. Thus late Paleozoic structures around the Antigonish Highlands reflect postcollisional, intracontinental deformation processes. The contrast in structural styles from local transtension (Stellarton basin; Yeo and Ruixiang, 1987; Waldron, 2004) to coeval transpression (Cape George Peninsula; St. Jean et al., 1993) and the deformation of the St. Mary’s Basin fill are consistent with dextral motion along major fault systems during this period. In addition, as noted by Gibling et al. (1992), this period was also notable for a change in paleocurrent directions, from those reflecting local derivation in the early Carboniferous to those reflecting regional derivation from the rising Appalachian foreland in the eastern United States by the late Carboniferous. Taken together, the Late Devonian to late Carboniferous tectonics around the Antigonish Highlands reflect episodic dextral shears that are collectively interpreted to accommodate local strains induced by the convergence and collision between
a
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ST. MARY'S BASIN AVALONIA MEGUMA
Figure 14. Schematic diagram showing the variation in style of the Horton Group (H)–Windsor Group (W) contact (after Murphy, 2003). In some sections, fault geometry has resulted in localized intense deformation and reverse faulting prior to Windsor Group deposition, such as is evident in the St. Mary’s basin (section a). Near the basin margins, motion along faults deformed the Horton Group strata prior to Windsor Group deposition, also resulting in an angular unconformity between the Horton and Windsor Groups (section b). In Horton Group sections that were distant from active faults, pre-Windsor deformation was minor, resulting in an approximately concordant contact (section c). The time gap implied by the fossil evidence indicates that this contact is a disconformity.
Gondwana and Laurussia. In contrast to the situation in the geologically complex regions of the northeastern United States, where the convergence and collision was more orthogonal and led to a complex interplay of deformation, metamorphism, and igneous intrusions (e.g., Hatcher, 2002), the manifestations of this event in this region of maritime Canada reflect oblique convergence and collision, so specific episodes in this progressive event are readily recognized in the stratigraphic record as well as in the sequence of deformational events. CONCLUSIONS Middle Ordovician–late Carboniferous strata in or adjacent to the Antigonish Highlands record both the regional tectonothermal events along the northern margin of the Rheic Ocean and those leading to the late Carboniferous amalgamation of Pangea. The Middle Ordovician Dunn Point Formation is thought to have formed in a microcontinental rifted arc setting similar to that of the Taupo volcanic zone in New Zealand. Deposition of the Arisaig Group is thought to have occurred after a soft collision of Avalonia with Baltica, and the lack of a pronounced angular unconformity between the Dunn Point Formation and the lowermost formation in the Arisaig Group suggests that this region was located along the trailing edge of Avalonia during terrane accretion. The geochemistry and isotopic signature of the Arisaig Group rocks clearly record the signal of these regional tectonic events. The accretion of Avalonia to Baltica by the Early Silurian and its accretion to Laurentia by the Middle Silurian are both constrained by these sedimentary sequences that overstep the terrane
Geological evolution of middle to late Paleozoic rocks in the Avalon terrane boundaries. Paleocontinental reconstructions indicate that the Early Silurian position of Western Avalonia lies within the range of detritus shed from coeval orogenic events that affected Baltica, whereas the Late Silurian–Early Devonian strata record loading of the margin, foreland basin development, and unroofing of more ancient (probably Laurentian) crust. Deformation of these strata in the Middle Devonian was accompanied by dextral shear along major transcurrent faults along the northern margin of the Rheic Ocean, which included between 5 and 12 km uplift of the Meguma terrane relative to the Avalon terrane. Late Devonian–early Carboniferous dextral shear along the boundary between the Avalon and Meguma terranes is recorded in ca. 370 Ma granites that occur along the southern boundary of the terrane, and continued dextral shear resulted in the formation of the St. Mary’s Basin. Middle to late Carboniferous strata record the far-field effects of the collision between Laurentia and Gondwana and the amalgamation of Pangea, which terminated the history of the Rheic Ocean. The style of tectonic activity varied from local transtension and basin formation to local transpression and intense thrusting and S-C fabric development. The Antigonish Highlands were located along the southern flank of the composite Maritimes basin, and sediment supply was profoundly influenced by mountain building associated with collisional tectonics. The stratigraphy, structure, and geochemistry of these rocks preserve evidence of local depocenter development, block uplifts, and basin inversion profoundly influenced by coeval dextral strike-slip on adjacent faults. ACKNOWLEDGMENTS My work has been supported by Natural Sciences and Engineering Research Council (Canada) Discovery, Research Capacity Development, and University Council for Research grants. I am grateful to Bernd Buschmann and Sue Johnson for very insightful and constructive reviews and to Jamie Braid, Mike Hamilton, Teresa Jeffries, Duncan Keppie, Randy Rice, Javier FernandezSuarez, and John Waldron for discussions. Contribution to International Geological Correlation Program Project 497. REFERENCES CITED Barr, S.M., White, C.E., and Miller, B.V., 2002, The Kingston terrane, southern New Brunswick, Canada: Evidence of an Early Silurian arc: Geological Society of America Bulletin, v. 114, p. 964–982, doi: 10.1130/00167606(2002)114<0964:TKTSNB>2.0.CO;2. Bergström, S.H., Huff, W.D., Kolata, D.R., and Melchin, M.J., 1997, Occurrence and significance of Silurian K-bentonite beds at Arisaig, Nova Scotia, eastern Canada: Canadian Journal of Earth Sciences, v. 34, p. 1630–1643. Bevier, M.L., Barr, S.M., and White, C.E., 1990, Late Precambrian U-Pb ages for the Brookville Gneiss, New Brunswick: Journal of Geology, v. 98, p. 955–968. Bevins, R.E., Bluck, B.E., and Brenchley, P.J., 1992, Ordovician, in Cope, J.C.W, Ingham, J.K., and Rawson, P.F., eds., Atlas of paleogeography and lithofacies, Geological Society of London Memoir 13, p. 19–36. Boehner, R.C., 1994, Carboniferous basin project activities 1993, in MacDonald, D., ed., Mines and Minerals Branch report of activities 1993: Halifax, Nova Scotia, Department of Natural Resources, Mines and Energy Branches Report 94–1, p. 123–130.
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Boehner, R.C., and Giles, P.S., 1982, Geological map of the Antigonish Basin, Nova Scotia: Halifax, Nova Scotia, Department of Mines and Energy Map 82–2, scale 1: 50,000. Boehner, R.C., and Giles, P.S., 1993, Geology of the Antigonish Basin, Antigonish County, Nova Scotia: Nova Scotia Department of Natural Resources, Mines and Minerals Branch Memoir 8, 109 p. Boucot, A.J., Dewey, J.F., Dineley, D.L., Fletcher, R., Fyson, W.K., Griffin, J.G., Hickox, C.F., McKerrow, W.S., and Ziegler, A.M., 1974, Geology of the Arisaig area, Antigonish County, Nova Scotia: Boulder, Colorado, Geological Society of America Special Paper 139, 191 p. Braid, J.A., and Murphy, J.B., 2006, Acadian deformation in the shallow crust: An example from the Siluro-Devonian Arisaig Group, Antigonish Highlands, Nova Scotia: Canadian Journal of Earth Sciences, v. 43, p. 157– 176, doi: 10.1139/e05-106. Cawood, P.A., McCausland, P.J.A., and Dunning, G.R., 2001, Opening Iapetus: Constraints from the Laurential margin of Newfoundland: Geological Society of America Bulletin, v. 113, p. 443–453, doi: 10.1130/00167606(2001)113<0443:OICFTL>2.0.CO;2. Chandler, F.W., 1997, Tectonics, paleoclimate and sandstone-hosted base metal deposits: Cause and effect in the Carboniferous Maritimes Basin, eastern Canada: Geological Association of Canada / Mineralogical Association of Canada Annual Meeting, Ottawa, v. 22, p. A24. Chandler, F.W., Waldron, J.W.F., Giles, P., and Gall, Q., 1997, Geology of the Stellarton Gap, Nova Scotia: Ottawa, Geological Society of Canada, Open File 3535, Map NTS 11 E/10 and parts of NTS 11/E7, 9, 15, scale 1:50,000. Clarke, D.B., MacDonald, M.A., and Tate, M.C., 1997, Late Devonian maficfelsic magmatism in the Meguma Zone, Nova Scotia, in Sinha, A.K., Whalen, J.B., and Hogan J.P., eds., Magmatism in the Appalachians: Boulder, Colorado, Geological Society of America Memoir 191, p. 107–127. Cocks, L.R.M., and Fortey, R.A., 1990, Biogeography of Ordovician and Silurian faunas, in McKerrow, W.S., and Scotese, C.R., eds., Paleozoic paleogeography and biogeography: Geological Society, London, Memoir 12, p. 97–104. Cocks, L.R.M., and Torsvik, T.H., 2002, Earth geography from 500 to 400 million years ago: A faunal and palaeomagnetic review: Journal of the Geological Society of London, v. 159, p. 631–644. Cocks, L.R.M., and Worsley, D., 1993, Late Llandovery and early Wenlock stratigraphy and ecology in the Oslo Region, Norway: Bulletin of the Natural History Museum, London, Geology Series, v. 49, p. 31–46. Cole, J.W., Darby, D.J., and Stern, T.A., 1995, Taupo Volcanic Zone and Central Volcanic Region: Backarc structures of North Island, New Zealand, in Taylor, B., ed., Backarc basins, tectonics and magmatism: New York and London, Plenum Press, p. 1–28. Cole, J.W., Gamble, J.A., Burt, R.M., Carroll, L.D., and Shelley, D., 2001, Mixing and mingling in the evolution of andesite-dacite magmas: Evidence from co-magmatic plutonic enclaves, Taupo Volcanic Zone, New Zealand: Lithos, v. 59, p. 25–46, doi: 10.1016/S0024-4937(01)00056-1. Connelly, J.N., and Heaman, L.M., 1993, U-Pb geochronological constraints on the tectonic evolution of the Grenville Province, western Labrador: Precambrian Research, v. 63, p. 123–142, doi: 10.1016/0301-9268(93)90008-P. Culshaw, N., and Liesa, M., 1997, Alleghanian reactivation of the Acadian Fold belt, Meguma Zone, southwest Nova Scotia: Canadian Journal of Earth Sciences, v. 34, p. 833–847. Dallmeyer, R.D., Keppie, J.D., and Nance, R.D., 1997, 40Ar/39Ar ages of detrital muscovite within Lower Cambrian and Carboniferous clastic sequences in northern Nova Scotia and southern New Brunswick: Implications for provenance regions: Canadian Journal of Earth Sciences, v. 34, p. 156–168. Daly, J.S., and McLelland, J.M., 1991, Juvenile middle Proterozoic crust in the Adirondack Highlands, Grenville Province, northeastern North America: Geology, v. 19, p. 119–122, doi: 10.1130/0091-7613(1991)019<0119: JMPCIT>2.3.CO;2. Dalziel, I.W.D., 1997, Overview: Neoproterozoic–Paleozoic geography and tectonics: Review, hypotheses and environmental speculations: Geological Society of America Bulletin, v. 109, p. 16–42, doi: 10.1130/00167606(1997)109<0016:ONPGAT>2.3.CO;2. Dalziel, I.A.W., Dalla Salda, L.H., and Gahagan, L.M., 1994, Paleozoic Laurentia-Gondwana interaction and the origin of the Appalachian-Andean mountain system: Geological Society of America Bulletin, v. 106, p. 243– 252, doi: 10.1130/0016-7606(1994)106<0243:PLGIAT>2.3.CO;2. Dickin, A.P., 2000, Crustal formation in the Grenville province: Nd isotopic evidence: Canadian Journal of Earth Sciences, v. 37, p. 165–181, doi: 10.1139/cjes-37-2-3-165.
432
Murphy
Dickin, A.P., and McNutt, R.H., 1989, Nd model age mapping of the southeast margin of the Archean foreland in the Grenville Province of Ontario: Geology, v. 17, p. 299–302, doi: 10.1130/0091-7613(1989)017<0299: NMAMOT>2.3.CO;2. Dickin, A.P., McNutt, R.H., and Clifford, P.M., 1990, A neodymium isotope study of plutons near the Grenville Front in Ontario, Canada: Chemical Geology, v. 83, p. 315–324, doi: 10.1016/0009-2541(90)90287-H. Dineley, D.L., 1963, Knoydart Formation (Lower Devonian) of Nova Scotia: Boulder, Colorado, Geological Society of America Special Paper 73, p. 138–139. Dunning, G.R., Barr, S.M., Giles, P.S., McGregor, D.C., Pe-Piper, G., and Piper, D.J.W., 2002, Chronology of Devonian to early Carboniferous rifting and igneous activity in southern Magdalen Basin based on U-Pb (zircon) dating: Canadian Journal of Earth Sciences, v. 39, p. 1219–1237, doi: 10.1139/e02-037. Durling, P., Harvey, P., and Howells, K., 1995a, Geophysical evidence for thrust faulting in the Carboniferous Antigonish-Mabou Subbasin, Nova Scotia: Atlantic Geology, v. 31, p. 183–196. Durling, P., Harvey, P., and Howells, K., 1995b, The near-surface geology of St. Georges Bay, Nova Scotia: Implications for the Hollow Fault: Canadian Journal of Earth Sciences, v. 32, p. 603–613. Fortey, R.A., and Cocks, L.R.M., 2003, Palaeontological evidence bearing on global Ordovician–Silurian continental reconstructions: Earth Science Reviews, v. 61, p. 245–307, doi: 10.1016/S0012-8252(02)00115-0. Fullagar, P.D., and Bottino, M.L., 1968, Radiometric age of the volcanic rocks at Arisaig, Nova Scotia, and the Ordovician-Silurian boundary: Canadian Journal of Earth Sciences, v. 5, p. 311–317. Gibling, M.R., Calder, J.H., Ryan, R., Van de Poll, H.W., and Yeo, G.M., 1992, Late Carboniferous and Early Permian drainage patterns in Atlantic Canada: Canadian Journal of Earth Sciences, v. 29, p. 338–352. Giles, P.S., 1981, Major transgressive-regressive cycles in the middle to Late Visean rocks of Nova Scotia: Halifax, Nova Scotia Department of Mines and Energy Paper 81–2, 27 p. Goldstein, S.J., and Jacobsen, S.B., 1988, Nd and Sr isotopic systematics of river water suspended material: Implications for crustal evolution: Earth and Planetary Science Letters, v. 87, p. 249–265, doi: 10.1016/0012821X(88)90013-1. Golonka, J., Ross, M.I., and Scotese, C.R., 1994, Phanerozoic paleogeographic and paleoclimatic modeling maps, in Beauchamp, B., Embry, A., and Glass, D., eds., Pangea: Calgary, Canadian Society of Petroleum Geologists Memoir 17, p. 1–47. Goodwin, A.M., 1991, Precambrian Geology: London, Academic Press, 666 p. Gower, C.F., and Tucker, R.D., 1994, The distribution of pre-1400 Ma crust in the Grenville province: Implications for rifting in Laurentia-Baltica during geon 14: Geology, v. 22, p. 827–830, doi: 10.1130/00917613(1994)022<0827:DOPMCI>2.3.CO;2. Gower, C.F., Ryan, A.B., and Rivers, T., 1990, Mid-Proterozoic Laurentia-Baltica: An overview of its geological evolution and a summary of the contributions made by this volume, in Gower, C.F., Rivers, T., and Ryan, A.B., eds. Mid-Proterozoic Laurentia-Baltica: St. John’s, Geological Association of Canada Special Paper 38, p.1–22. Gutierrez-Alonso, G., Fernandez-Suarez, J., Collins, A.S., Abad, I., and Nieto, F., 2005, An Amazonian Mesoproterozoic basement in the core of the Ibero-Armorican Arc: 40Ar/39Ar detrital mica ages complement the zircon’s tale: Geology, v. 33, p. 637–640, doi: 10.1130/G21485.1. Hamblin, A.P., and Rust, B.R., 1989, Tectono-sedimentary analysis of alternate-polarity half-graben basin-fill successions: The Late Devonian–Early Carboniferous Horton Group, Cape Breton Island, Nova Scotia, Canada: Basin Research, v. 2, p. 239–255. Hamilton, M.A., and Murphy, J.B., 2004, Tectonic significance of a Llanvirn age for the Dunn Point volcanic rocks, Avalon terrane, Nova Scotia, Canada: Implications for the evolution of the Iapetus and Rheic oceans: Tectonophysics, v. 379, p. 199–209, doi: 10.1016/j.tecto.2003.11.006. Hatcher, R.D., Jr., 2002, The Alleghanian (Appalachian) orogeny, a product of zipper tectonics: Rotational transpressive continent-continent collision and closing of ancient oceans along irregular margins, in Catalán, J.R.M., Hatcher, R.D., Jr., Arenas, R., and García, F.D., eds., Variscan-Appalachian dynamics: The building of the late Paleozoic basement: Boulder, Colorado, Geological Society of America Special Paper 364, p. 199–208. Hicks, R., Jamieson, R., and Reynolds, P.H., 1999, Detrital and metamorphic 40 Ar/39Ar ages from muscovite and whole-rock samples, Meguma Supergroup, southern Nova Scotia: Canadian Journal of Earth Sciences, v. 36, p. 23–32, doi: 10.1139/cjes-36-1-23.
Hodych, J.P., and Buchan, K.L., 1998, Palaeomagnetism of the ca. 440 Ma Cape St Mary’s sills of the Avalon Peninsula of Newfoundland: Implications for Iapetus Ocean closure: Geophysical Journal International, v. 135, p. 155–164, doi: 10.1046/j.1365-246X.1998.00263.x. Hoffman, P.F., 1989, Precambrian geology and tectonic history of North America, in Bally, A.M., and Palmer, A.R., eds. The geology of North America: An overview, v. A, Geology of North America: Boulder, Colorado, Geological Society of America, p. 447–512. Hurst, J.M., and Pickerill, R.K., 1986, The relationship between sedimentary facies and faunal associations in the Llandovery siliciclastics Ross Brook Formation, Arisaig, Antigonish County, Nova Scotia: Canadian Journal of Earth Sciences, v. 23, p. 711–716. Jennex, L.C., Murphy, J.B., and Anderson, A.J., 2000, Post-orogenic exhumation of an auriferous terrane: The paleoplacer potential of the early Carboniferous St. Marys Basin, Canadian Appalachians: Mineralium Deposita, v. 35, p. 776–790, doi: 10.1007/s001260050279. Johnson, R.J.E., and Van der Voo, R., 1986, Paleomagnetism of the Late Precambrian Fourchu Group, Cape Breton Island, Nova Scotia: Canadian Journal of Earth Sciences, v. 23, p. 1673–1685. Johnson, R.J.E., and Van der Voo, R., 1990, Pre-folding magnetization reconfirmed for the Late Ordovician–Early Silurian Dunn Point volcanics, Nova Scotia: Tectonophysics, v. 178, p. 193–205, doi: 10.1016/00401951(90)90146-Y. Karabinos, P., and Gromet, L.P., 1993, Application of single-grain zircon evaporation analyses to detrital grain studies and age discrimination in igneous suites: Geochimica et Cosmochimica Acta, v. 57, p. 4257–4267, doi: 10.1016/0016-7037(93)90321-M. Keppie, J.D., 1985. The Appalachian Collage, in Gee, D.G., and Sturt, B., eds., The Caledonide Orogen, Scandinavia, and related areas: New York, J. Wiley and Sons, p. 1217–1226. Keppie, J.D., and Dallmeyer, R.D., 1995, Late Paleozoic collision, delamination, short-lived magmatism and rapid denudation in the Meguma terrane (Nova Scotia, Canada): Constraints from 40Ar/ 39Ar isotopic data: Canadian Journal of Earth Sciences, v. 32, p. 644–659. Keppie, J.D., and Smith, P.K., 1978, Compilations of isotopic data of Nova Scotia: Halifax, Nova Scotia, Department of Mines Report 78–4. Keppie, J.D., Giles, P.S., and Boehner, R.C., 1978, Some Middle Devonian to Lower Devonian rocks of Cape George, Nova Scotia: Halifax, Nova Scotia, Department of Mines Report 78–4. Keppie, J.D., Dostal, J., and Zentilli, M., 1979, Early Silurian volcanic rocks at Arisaig, Nova Scotia: Canadian Journal of Earth Sciences, v. 16, p. 1635–1639. Keppie, J.D., Dostal, J., Murphy, J.B., and Cousens, B.L., 1997, Palaeozoic within-plate volcanic rocks in Nova Scotia (Canada) reinterpreted: Isotopic constraints on magmatic source and paleocontinental reconstructions: Geological Magazine, v. 134, p. 425–447, doi: 10.1017/ S001675689700719X. Keppie, J.D., Davis, D.W., and Krogh, T.E., 1998, U-Pb geochronological constraints on Precambrian stratified units in the Avalon composite terrane of Nova Scotia, Canada: Tectonic implications: Canadian Journal of Earth Sciences, v. 35, p. 222–236, doi: 10.1139/cjes-35-3-222. Keppie, J.D., Nance, R.D., Murphy, J.B., and Dostal, J., 2003, Tethyan, Mediterranean, and Pacific analogues for the Neoproterozoic–Paleozoic birth and development of peri-Gondwanan terranes and their transfer to Laurentia and Laurussia: Tectonophysics, v. 365, p. 195–220, doi: 10.1016/ S0040-1951(03)00037-4. Krogh, T.E., and Keppie, J.D., 1990, Age of detrital zircon and titanite in the Meguma Group, southern Nova Scotia, Canada: Clues to the origin of the Meguma Terrane: Tectonophysics, v. 177, p. 307–323, doi: 10.1016/00401951(90)90287-I. Landing, E., 1996, Avalon: Insular continent by the latest Precambrian, 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. 29–63. Landing, E., 2005, Early Paleozoic Avalon-Gondwana unity: An obituaryresponse to “Palaeontological evidence bearing on global Ordovician– Silurian continental reconstructions” by R.A. Fortey and L.R.M. Cocks: Discussion: Earth-Science Reviews, v. 69, p. 169–175, doi: 10.1016/ j.earscirev.2004.10.002. Landing, E., and Murphy, J.B., 1991, Uppermost Precambrian (?)–Lower Cambrian of mainland Nova Scotia: Faunas, depositional environments and stratigraphic revision: Journal of Paleontology, v. 65, p. 382–396.
Geological evolution of middle to late Paleozoic rocks in the Avalon terrane Lane, T.E., and Jensen, L.R., 1975, Stratigraphy of the Arisaig Group: Maritime Sediments, v. 11, p. 119–140. Lees, D., Fortey, R.A., and Cocks, L.R.M., 2002, Quantifying paleogeography using biogeography: A test case for the Ordovician and Silurian of Avalonia based on brachiopods and trilobites: Paleobiology, v. 28, p. 343–363, doi: 10.1666/0094-8373(2002)028<0343:QPUBAT>2.0.CO;2. Linnemann, U., Gehmlich, M., Tichomirowa, M., Buschmann, B., Nasdala, L., Jonas, P., Lützner, H., and Bombach, K., 2000, From Cadomian subduction to Early Palaeozoic rifting: The evolution of Saxo-Thuringia at the margin of Gondwana in the light of single zircon geochronology and basin development (Central European Variscides, Germany), in Franke, W., Altherr, R., Haak, V., Oncken, O., and Tanner, D., eds., Orogenic processes: Quantification and modelling in the Variscan Belt of Central Europe: Geological Society of London Special Publication 179, p. 131–153. MacLachlan, K., and Dunning, G., 1998, U-Pb ages and tectonomagmatic relationships of early Ordovician low-Ti tholeiites, boninites, and related plutonic rocks in central Newfoundland: Contributions to Mineralogy and Petrology, v. 133, p. 235–258, doi: 10.1007/s004100050450. MacNiocaill, C., 2000, A new Silurian paleolatitude for eastern Avalonia and evidence for crustal rotations in the Avalonian margin of southeast Ireland: Geophysical Journal International, v. 141, p. 661–671, doi: 10.1046/ j.1365-246x.2000.00101.x. MacNiocaill, C., and Smethurst, M.A., 1994, Palaeozoic palaeogeography of Laurentia and its margins: A reassessment of the paleomagnetic data: Geophysical Journal International, v. 116, p. 715–725. MacNiocaill, C., van der Pluijm, B.A., and Van der Voo, R., 1997, Ordovician paleogeography and the evolution of the Iapetus Ocean: Geology, v. 25, p. 159–162, doi: 10.1130/0091-7613(1997)025<0159:OPATEO>2.3.CO;2. Marillier, F., Keen, C.E., Stockmal, G.S., Quinlan, G., Williams, H., ColmanSadd, S.P., and O’Brien, S.J., 1989, Crustal structure and surface zonation of the Canadian Appalachians: Implications for deep seismic reflection data: Canadian Journal of Earth Sciences, v. 26, p. 305–321. Martel, A.T., 1987, Seismic stratigraphy and hydrocarbon potential of the strikeslip Sackville sub-basin, New Brunswick, in Beaumont, C., and Tankard, A.J., eds., Sedimentary basins and basin-forming mechanisms: Calgary, Canadian Society of Petroleum Geologists Memoir 12, p. 363–377. Martel, A.T., and Gibling, M.R., 1996, Stratigraphy and tectonic history of the Upper Devonian–Lower Carboniferous Horton Bluff Formation, Nova Scotia: Atlantic Geology, v. 32, p. 13–38. Martel, A.T., McGregor, D.C., and Utting, J., 1993, Stratigraphic significance of Upper Devonian and Lower Carboniferous miospores from the Horton Group type area, Nova Scotia: Canadian Journal of Earth Sciences, v. 30, p. 1091–1098. Martignole, J., Machado, N., and Indares, A., 1994, The Wakeham terrane: A Mesoproterozoic terrestrial rift in the eastern part of the Grenville province: Precambrian Research, v. 68, p. 291–306, doi: 10.1016/03019268(94)90034-5. McCutcheon, S.R., and Robinson, P.T., 1987, Geological constraints on the genesis of the Maritimes basin, Atlantic Canada, in Beaumont, C., and Tankard, A.J., eds., Sedimentary basins and basin-forming mechanisms: Calgary, Canadian Society of Petroleum Geologists Memoir 12, p. 287–297. McKerrow, W.S., and Scotese, C.R., eds., 1990, Palaeozoic palaeogeography and biogeography: Geological Society of London Memoir 12, 435 p. McKerrow, W.S., Dewey, J.F., and Scotese, C.R., 1991, The Ordovician and Silurian development of the Iapetus Ocean: Special Papers in Paleontology, v. 44, p. 165–178. McLelland, J.M., Daly, J.S., and Chiarenzelli, J., 1993, Sm-Nd and U-Pb isotopic evidence of juvenile crust in the Adirondack lowlands and implications for the evolution of the Adirondack Mountains: Journal of Geology, v. 101, p. 97–105. McNamara, A.K., Mac Niocaill, C., van der Pluijm, B.A., and Van der Voo, R., 2001, West African proximity of the Avalon terrane in the latest Precambrian: Geological Society of America Bulletin, v. 113, p. 1161–1170, doi: 10.1130/0016-7606(2001)113<1161:WAPOTA>2.0.CO;2. Miller, B.V., Nance, R.D., and Murphy, J.B., 1995, Kinematics of the Rockland Brook Fault, Nova Scotia: Implications for the interaction of the Meguma and Avalon terranes: Journal of Geodynamics, v. 19, p. 253–270, doi: 10.1016/0264-3707(94)00017-P. Miller, J.D., and Kent, D.V., 1988, Paleomagnetism of the Siluro-Devonian Andreas redbeds: Evidence of a Devonian supercontinent?: Geology, v. 16, p. 195–198, doi: 10.1130/0091-7613(1988)016<0195:POTSDA> 2.3.CO;2.
433
Miller, R.G., O’Nions, R.K., Hamilton, P.J., and Welin, E., 1986, Crustal residence ages of clastic sediments, orogeny and crustal evolution: Chemical Geology, v. 57, p. 87–99, doi: 10.1016/0009-2541(86)90095-1. Murphy, J.B., 1987, The stratigraphy and depositional environment of upper Ordovician to lower Devonian rocks in the Antigonish Highlands, Nova Scotia: Maritime Sediments and Atlantic Geology, v. 23, p. 63–75. Murphy, J.B., 2000, Tectonic influence on sedimentation along the southern flank of the Late Paleozoic Magdalen Basin in the Canadian Appalachians: Geochemical and isotopic constraints on the Horton Group in the St. Mary’s Basin, Nova Scotia: Geological Society of America Bulletin, v. 112, p. 997–1011, doi: 10.1130/0016-7606(2000)112<0997:TIOSAT>2.3.CO;2. Murphy, J.B., 2001, Geochemistry and tectonic significance of the middle to late Devonian McAras Brook and the Visean Martin Road Formation, Merigomish sub-basin, Nova Scotia: Atlantic Geology, v. 37, p. 161–173. Murphy, J.B., 2003, Late Paleozoic formation and development of the St. Marys Basin, mainland Nova Scotia, Canada: A prolonged record of intra-continental strike-slip deformation during the assembly of Pangaea, in Storti, F., Holdsworth, R.E., and Salvini, F., eds., Intraplate strike-slip deformation belts: Geological Society of London Special Publication 210, p. 185–196. Murphy, J.B., and Hamilton, M.A., 2000, U-Pb detrital zircon age constraints on evolution of the Late Paleozoic St. Marys Basin, central mainland Nova Scotia: Journal of Geology, v. 108, p. 53–72, doi: 10.1086/314384. Murphy, J.B., and Keppie, J.D., 1998, Late Devonian palinspastic reconstruction of the Avalon-Meguma terrane boundary: Implications for terrane accretion and basin development in the Appalachian orogen: Tectonophysics, v. 284, p. 221–231, doi: 10.1016/S0040-1951(97)00177-7. Murphy, J.B., and Keppie, J.D., 2005, The Acadian orogeny in the northern Appalachians: International Geology Review, v. 47, p. 663–687. Murphy, J.B., and MacDonald, D.A., 1993, Geochemistry of Late Proterozoic arc-related volcaniclastic turbidite sequences, Antigonish Highlands, Nova Scotia: Canadian Journal of Earth Sciences, v. 30, p. 2273–2282. Murphy, J.B., and Nance, R.D., 1989, A model for the evolution of the Avalonian-Cadomian belt: Geology, v. 17, p. 735–738, doi: 10.1130/00917613(1989)017<0735:MFTEOT>2.3.CO;2. Murphy, J.B., and Nance, R.D., 1991, A supercontinent model for the contrasting character of Late Proterozoic orogenic belts: Geology, v. 19, p. 469– 472, doi: 10.1130/0091-7613(1991)019<0469:SMFTCC>2.3.CO;2. Murphy, J.B., and Nance, R.D., 2002, Nd-Sm isotopic systematics as tectonic tracers: An example from West Avalonia, Canadian Appalachians: Earth Science Reviews, v. 59, p. 77–100, doi: 10.1016/S0012-8252(02)00070-3. Murphy, J.B., and Rice, R.J., 1998, Stratigraphy and depositional environment of Horton Group rocks in the St. Mary’s Basin, central mainland Nova Scotia: Atlantic Geology, v. 34, p. 1–26. Murphy, J.B., Cameron, K., Dostal, J., Keppie, J.D., and Hynes, A.J., 1985, Cambrian volcanism in Nova Scotia, Canada: Canadian Journal of Earth Sciences, v. 22, p. 599–606. Murphy, J.B., Keppie, J.D., and Hynes, A.J., 1991, Geology of the Antigonish Highlands: Ottawa, Geological Survey of Canada Paper 89–10, 114 p. Murphy, J.B., Nance, R.D., Keppie, J.D., Dostal, J., and Cousens, B.L., 1995a, The Odyssey of West Avalonia: Isotopic signatures as constraints for the Late Proterozoic–Early Silurian paleogeography, in Hibbard, J., van Staal, C., and Cawood, P., eds., New perspectives in the Caledonian-Appalachian orogen: St. John’s, Geological Association of Canada Special Paper 41, p. 227–238. Murphy, J.B., Rice, R.J., Stokes, T.R., and Keppie, D.F., 1995b, The St. Mary’s Basin, central mainland Nova Scotia: Late Paleozoic basin formation and deformation along the Avalon-Meguma Terrane boundary, Canadian Appalachians, in Hibbard, J., Van Staal, C., and Cawood, P., eds., New perspectives in the Caledonian-Appalachian orogen: St. John’s, Geological Association of Canada Special Paper 41, p. 409–420. Murphy, J.B., Keppie, J.D., Dostal, J., Waldron, J.W.F., and Cude, M.P., 1996a, Geochemical and isotopic characteristics of Early Silurian clastic sequences in Antigonish Highlands, Nova Scotia, Canada: Constraints on the accretion of Avalonia in the Appalachian- Caledonide Orogen: Canadian Journal of Earth Sciences, v. 33, p. 379–388. Murphy, J.B., Keppie, J.D., Dostal, J., and Cousins, B.L., 1996b, Repeated late Neoproterozoic–Silurian lower crustal melting beneath the Antigonish Highlands, Nova Scotia: Nd isotopic evidence and tectonic interpretations, in Nance, R.D., and Thompson, M.D., eds., Avalonian and related peri-Gondwanan terranes of the circum–North Atlantic: Boulder, Colorado, Geological Society of America Special Paper 304, p. 109–120.
434
Murphy
Murphy, J.B., van Staal, C.R., and Keppie, J.D., 1999a, Is the mid to late Paleozoic Acadian Orogeny a plume-modified Laramide-style orogeny?: Geology, v. 27, p. 653–656, doi: 10.1130/0091-7613(1999)027<0653:MTLPAO>2.3.CO;2. Murphy, J.B., Keppie, J.D., Dostal, J., and Nance, R.D., 1999b, Neoproterozoic–Early Paleozoic evolution of Avalonia and Cenozoic analogues, in Ramos, V., and Keppie, J.D., eds., Laurentia-Gondwana connections before Pangea: Boulder, Colorado, Geological Society of America Special Paper 336, p. 253–266. Murphy, J.B., Strachan, R.A., Nance, R.D., Parker, K.D., and Fowler, M.B., 2000, Proto-Avalonia: A 1.2 to 1.0 Ga tectonothermal event and constraints for the evolution of Rodinia: Geology, v. 28, p. 1071–1074, doi: 10.1130/00917613(2000)028<1071:PAAGTE>2.3.CO;2. Murphy, J.B., Fernández-Suárez, J., Jeffries, T.E., and Strachan, R.A., 2004b, LA-ICP-MS U-Pb detrital zircon data from Cambrian clastic rocks in Avalonia: Erosion of a Neoproterozoic arc along the northern Gondwanan margin: Journal of the Geological Society, London, v. 161, p. 243–254. Murphy, J.B., Fernández-Suárez, J., and Jeffries, T.E., 2004c, Lithogeochemical, Sm-Nd and U-Pb isotopic data from the Silurian–Early Devonian Arisaig Group clastic rocks, Avalon terrane, Nova Scotia: A record of terrane accretion in the Appalachian-Caledonide orogen: Geological Society of America Bulletin, v. 116, p. 1183–1201, doi: 10.1130/B25423.1. Murphy, J.B., Fernández-Suárez, J., Keppie, J.D., and Jeffries, T.E., 2004a, Contiguous rather than discrete Paleozoic histories for the Avalon and Meguma terranes based on detrital zircon data: Geology, v. 32, p. 585– 588, doi: 10.1130/G20351.1. Murphy, J.B., Pisarevsky, S.A., Nance, R.D., and Keppie, J.D., 2004d, Neoproterozoic–early Paleozoic configuration of peri-Gondwanan terranes: Implications for Laurentia-Gondwanan connections: International Journal of Earth Sciences, v. 93, p. 659–682, doi: 10.1007/s00531-004-0412-9. Nance, R.D., and Murphy, J.B., 1994, Contrasting basement isotopic signatures and the palinspastic restoration of peripheral orogens: Example from the Neoproterozoic Avalonian-Cadomian belt: Geology, v. 22, p. 617–620, doi: 10.1130/0091-7613(1994)022<0617:CBISAT>2.3.CO;2. Nance, R.D., and Murphy, J.B., 1996, Basement isotopic signatures and Neoproterozoic paleogeography of Avalonian-Cadomian and related terranes in the circum North Atlantic, in Nance, R.D., and Thompson, M.D., eds., Avalonian and related peri-Gondwanan terranes of the circum–North Atlantic: Boulder, Colorado, Geological Society of America Special Paper 304, p. 333–346. Nance, R.D., and Thompson, M.D., 1996, eds., Avalonian and related periGondwanan terranes of the circum–North Atlantic: Boulder, Colorado, Geological Society of America Special Paper 304, 390 p. O’Brien, S.J., Wardle, R.J., and King, A.F., 1983, The Avalon zone: A PanAfrican terrane in the Appalachian orogen of Canada: Geological Journal, v. 18, p. 195–222. Patchett, P.J., and Ruiz, J., 1989, Nd isotopes and the origin of the Grenville-age rocks in Texas: Implications for Proterozoic evolution of the United States mid-continent region: Journal of Geology, v. 97, p. 685–695. Pedersen, R.B., and Dunning, G.R., 1997, Evolution of arc crust and relations between contrasting sources: U-Pb (age), Nd and Sr isotopic systematics of the ophiolite terrain of SW Norway: Contributions to Mineralogy and Petrology, v. 128, p. 1–15, doi: 10.1007/s004100050289. Pickerill, R.K., and Hurst, J.M., 1983, Sedimentary facies, depositional environments, and faunal associations of the lower Llandovery (Silurian) Beechill Cove Formation: Canadian Journal of Earth Sciences, v. 20, p. 1761–1779. Pickering, K., Basset, M.G., and Siveter, D.J., 1988, Late Ordovician–Early Silurian destruction of the Iapetus Ocean: Newfoundland, British Isles and Scandinavia—A discussion: Transactions of the Royal Society of Edinburgh, v. 79, p. 361–382. Piper, D.J.W., and Pe-Piper, G., 2001, Tectonic deformation and magmatism along the southern flank of the Maritimes Basin: The northeastern Cobequid Highlands, Nova Scotia: Canadian Journal of Earth Sciences, v. 38, p. 43–58, doi: 10.1139/cjes-38-1-43. Potts, S., van der Pluijm, B., and Van der Voo, R., 1993, Discordant Silurian paleolatitudes for central Newfoundland: New paleomagnetic evidence from the Springdale Group: Earth and Planetary Science Letters, v. 120, p. 1–12, doi: 10.1016/0012-821X(93)90019-6. Roberts, D., 2003, The Scandinavian Caledonides: Event chronology, palaeogeographic settings and likely modern analogues: Tectonophysics, v. 365, p. 283–299, doi: 10.1016/S0040-1951(03)00026-X. Roberts, D., and Siedlecka, A., 2002, Timanian orogenic deformation along the northeastern margin of Baltica, Northwest Russia and Northeast Norway,
and Avalonian-Cadomian connections: Tectonophysics, v. 352, p. 169– 184, doi: 10.1016/S0040-1951(02)00195-6. Rocci, G., Bronner, G., and Deschamps, M., 1991, Crystalline basement of the West African craton, in Dallmeyer, R.D., and Lecorche, J.P., eds., The West African orogens and circum-Atlantic correlatives: Heidelberg, Germany, Springer, p. 31–61. Ryan, R.J., and Boehner, R.C., 1994, Geology of the Cumberland Basin, Cumberland, Colchester, and Pictou Counties, Nova Scotia: Halifax, Nova Scotia, Department of Natural Resources, Mines and Energy Branches, Memoir 10, 222 p. Ryan, R.J., Calder, J.H., Donohoe, H.V., Jr., and Naylor, R., 1987, Late Paleozoic sedimentation and basin development adjacent to the Cobequid Highlands Massif, Eastern Canada, in Beaumont, C., and Tankard, A.J., eds., Sedimentary basins and basin-forming mechanisms: Calgary, Canadian Society of Petroleum Geologists Memoir 12, p. 311–317. Sadowski, G.R., and Bettencourt, J.S., 1996, Mesoproterozoic tectonic correlations between eastern Laurentia and the western border of the Amazon craton: Precambrian Research, v. 76, p. 213–227, doi: 10.1016/03019268(95)00026-7. Schenk, P.E., 1997, Sequence stratigraphy and provenance on Gondwana’s margin: The Meguma zone (Cambrian–Devonian) of Nova Scotia, Canada: Geological Society of America Bulletin, v. 109, p. 395–409, doi: 10.1130/0016-7606(1997)109<0395:SSAPOG>2.3.CO;2. Soper, N.J., and Woodcock, N.H., 1990, Silurian collision and sediment dispersal patterns in southern Britain: Geological Magazine, v. 127, p. 527–542. St. Jean, J.A.R., Nance, R.D., and Murphy, J.B., 1993, Tectonic significance of Late Paleozoic deformation, Cape George Peninsula, Nova Scotia: Atlantic Geology, v. 29, p. 27–42. St. Peter, C., 1993, Maritime basin evolution: Key geologic and seismic evidence from the Moncton Subbasin of New Brunswick: Atlantic Geology, v. 29, p. 233–270. Stampfli, G.M., and Borel, G.D., 2002, A plate tectonic model for the Paleozoic and Mesozoic constrained by dynamic plate boundaries and restored synthetic oceanic isochrones: Earth and Planetary Science Letters, v. 196, p. 17–33, doi: 10.1016/S0012-821X(01)00588-X. Starmer, I.C., 1993, The Sveconorwegian Orogeny in southern Norway, relative to deep crustal structures and events in the North Atlantic Proterozoic supercontinent: Norsk Geologisk Tidsskrift, v. 73, p. 109–132. Stevens, J.E., Murphy, J.B., and Chandler, F.W., 1999, Geochemistry of the Namurian Lismore Formation, northern mainland Nova Scotia: Sedimentation and tectonic activity along the southern flank of the Maritimes Basin: Canadian Journal of Earth Sciences, v. 36, p. 1655–1669, doi: 10.1139/cjes-36-10-1655. Strachan, R.A., and Taylor, G.K., eds., 1990, Avalonian and Cadomian geology of the North Atlantic: London, Blackie, 252 p. Teixiera, W., Tassinari, C.C.G., Cordani, U.G., and Kawashita, K., 1989, A review of the geochronology of the Amazon craton: Tectonic implications: Precambrian Research, v. 42, p. 213–227, doi: 10.1016/0301-9268(89)90012-0. Theokritoff, G., 1979, Early Cambrian provincialism and biogeographic boundaries in the North Atlantic region: Lethaia, v. 12, p. 281–295. Thomas, D., Nance, R.D., and Murphy, J.B., 2002, Deformation of the Macumber Formation, Antigonish Basin, Nova Scotia: Implications for the Ainslie detachment: Atlantic Geology, v. 38, p. 135–144. Thorogood, E.J., 1990, Provenance of the pre-Devonian sediments of England and Wales: Sm-Nd isotopic evidence: Journal of the Geological Society, London, v. 147, p. 591–594. Torsvik, T.H., and Rehnström, E.F., 2003, The Tornquist Sea and BalticaAvalonia docking: Tectonophysics, v. 362, p. 67–82, doi: 10.1016/S00401951(02)00631-5. Trench, A., and Torsvik, T.H., 1992, The closure of the Iapetus Ocean and Tornquist Sea: New paleomagnetic constraints: Journal of the Geological Society of London, v. 149, p. 867–870. Utting, J., 1987, Palynology of the Lower Carbonifeorus Windsor Group and Windsor-Canso Boundary Beds of Nova Scotia, and their equivalents in Quebec, New Brunswick and Newfoundland: Utting, Geological Survey of Canada Bulletin 374, 93 p. Utting, J., and Hamblin, A.P., 1991, Thermal maturity of the Lower Carboniferous Horton Group, Nova Scotia: International Journal of Coal Geology, v. 19, p. 439–456, doi: 10.1016/0166-5162(91)90029-I. Van der Voo, R., 1988, Palaeozoic paleogeography of North America, Gondwana, and intervening displaced terranes: Comparisons of palaeomagnetism within palaeoclimatology and biogeographical patterns: Geological
Geological evolution of middle to late Paleozoic rocks in the Avalon terrane Society of America Bulletin, v. 100, p. 311–324, doi: 10.1130/00167606(1988)100<0311:PPONAG>2.3.CO;2. Van der Voo, R., and Johnson, R.J.E., 1985, Paleomagnetism of the Dunn Point Formation (Nova Scotia): High paleolatitudes for the Avalon terrane in the Late Ordovician: Geophysical Research Letters, v. 12, p. 337–340. van Staal, C.R., Ravenhurst, C.E., Winchester, J.A., Roddick, J.C., and Langton, J.P., 1990, Post Taconic blueschist suture in the northern Appalachians of New Brunswick, Canada: Geology, v. 18, p. 1073–1077. van Staal, C.R., Sullivan, R.W., and Whalen, J.B., 1996, Provenance and tectonic history of the Gander Zone in the Caledonide/Appalachian orogen: Implications for the origin and assembly of Avalon, in Nance, R.D., and Thompson, M.D., eds., Avalonian and related peri-Gondwanan terranes of the circum–North Atlantic: Boulder, Colorado, Geological Society of America Special Paper 304, p. 347–367. van Staal, C.R., Dewey, J.F., Mac Niocaill, C., and McKerrow, W.S., 1998, The Cambrian–Silurian tectonic evolution of the Northern Appalachians and British Caledonides: History of a complex, west and southwest Pacifictype segment of Iapetus, in Blundell, D., and Scott, A.C., eds., Lyell: The past is the key to the present: Geological Society of London Special Publication 143, p. 199–242. Waldron, J.W.F., 2004, Anatomy and evolution of the Stellarton pull-apart basin, Nova Scotia: Geological Society of America Bulletin, v. 116, p. 109–127, doi: 10.1130/B25312.1. Waldron, J.W.F., Murphy, J.B., Melchin, M., and Davis, G., 1996, Silurian tectonics of western Avalonia: Strain corrected subsidence history of the Arisaig Group, Nova Scotia: Journal of Geology, v. 104, p. 677–694.
435
Williams, H., 1979, The Appalachian orogen in Canada: Canadian Journal of Earth Sciences, v. 16, p. 792–807. Williams, H., and Hatcher, R.D., 1983, Appalachian suspect terranes, in Hatcher, R.D., Williams, H., and Zietz, I., eds., Contributions to the tectonics and geophysics of mountain chains: Boulder, Colorado, Geological Society of America Memoir 158, p. 33–53. Williams, M., Stone, P., Siveter, D.J., and Taylor, P., 2001, Upper Ordovician ostracods from the Cautley district, northern England: Baltic and Laurentian affinities: Geological Magazine, v. 138, p. 599–607. Williams, S.H., Harper, D.A.T., Neuman, R.B., Boyce, W.D., and Mac Niocaill, C., 1995, Lower Paleozoic fossils from Newfoundland and their importance in understanding the history of the Iapetus Ocean, in Hibbard, J., van Staal, C., and Cawood, P., eds., Current perspectives in the Appalachian-Caledonian orogen: St. John’s, Geological Association of Canada Special Paper 41, p. 115–126. Woodcock, N., and Strachan, R.A., 2000, Geological history of the British Isles: London, Blackwell Science. Yeo, G.M., and Ruixiang, G., 1987, Stellarton Graben: An upper Carboniferous pull-apart basin in northern Nova Scotia, in Beaumont, C., and Tankard, A.J., eds., Sedimentary basins and basin-forming mechanisms: Calgary, Canadian Society of Petroleum Geologists Memoir 12, p. 299–309.
MANUSCRIPT ACCEPTED BY THE SOCIETY 3 OCTOBER 2006
Printed in the USA
Geological Society of America Special Paper 423 2007
Vestige of the Rheic Ocean in North America: The Acatlán Complex of southern México R. Damian Nance* Department of Geological Sciences, 316 Clippinger Laboratories, Ohio University, Athens, Ohio 45701, USA Brent V. Miller Department of Geology and Geophysics, Texas A&M University, College Station, Texas 77843, USA J. Duncan Keppie Departamento de Geología Regional, Instituto de Geología, Universidad Nacional Autónoma de México, 04510 México D.F., México J. Brendan Murphy Department of Earth Sciences, St. Francis Xavier University, Antigonish, Nova Scotia, Canada, B2G 2W5 Jaroslav Dostal Department of Geology, St. Mary’s University, Halifax, Nova Scotia, Canada, B3H 3C3
ABSTRACT The Acatlán Complex of southern México comprises metasedimentary and metaigneous rocks that represent the vestige of a Paleozoic ocean. Juxtaposed against granulite-facies gneisses of Mesoproterozoic (ca. 1 Ga) age, the complex has previously been related to the Iapetus Ocean and interpreted to preserve a tectonostratigraphic record linked to that of the Appalachian orogen: (1) Cambro-Ordovician deposition of a trench or forearc sequence (the Petlalcingo Group: the Magdalena, Chazumba, and Cosoltepec Formations) and an oceanic assemblage (the Piaxtla Group), (2) polyphase Late Ordovician–Early Silurian deformation (the Acatecan orogeny) during which the Piaxtla Group underwent eclogite-facies metamorphism synchronous with megacrystic granitoid emplacement, (3) deposition of the arc-related Tecomate Formation and intrusion of megacrystic granitoid plutons during the Devonian, and (4) deformation under greenschist-facies conditions during the Late Devonian Mixtecan orogeny. However, recent structural, geochronological, and geochemical studies have shown that (1) the Cosoltepec Formation is bracketed between ca. 455 Ma and the latest Devonian and may be part of a continental rise prism with slivers of oceanic basalt; (2) the Magdalena and Chazumba Units represent a clastic wedge assemblage
*E-mail:
[email protected]. Nance, R.D., Miller, B.V., Keppie, J.D., Murphy, J.B., and Dostal, J., 2007, Vestige of the Rheic Ocean in North America: The Acatlán Complex of southern México, 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. 437–452, doi: 10.1130/2007.2423(21). For permission to copy, contact
[email protected]. ©2007 Geological Society of America. All rights reserved.
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Nance et al. of Permo-Triassic age; (3) the eclogitic metamorphism is locally Mississippian in age; (4) the Tecomate Formation is an arc complex of latest Pennsylvanian–Middle Permian age; (5) the megacrystic granitoid rocks span the Ordovician and have a calc-alkaline geochemistry, whereas accompanying mafic units have mixed continental arc–tholeiitic affinities and are locally as young as the earliest Silurian; (6) the greenschist-facies tectonothermal event occurred in the Permo-Triassic; and (7) the complex records a Jurassic tectonothermal event that resulted in local high-grade metamorphism and migmatization. This revised geological history precludes any linkage to Iapetus, but is consistent with that of the Rheic and paleo-Pacific Oceans and is interpreted to record (1) development of a rift or passive margin on the southern flank of the Rheic Ocean in the Cambro-Ordovician, (2) formation of either an arc or an extensional regime along the formerly active northern margin of Gondwana throughout the Ordovician, (3) ocean closure documented by subduction-related eclogite-facies metamorphism and exhumation during the Late Devonian–Mississippian, (4) Permo-Triassic convergent tectonics on the paleo-Pacific margin of Pangea, and (5) overriding of a Jurassic plume. Keywords: Acatlán Complex, geochronology, Paleozoic, México, Rheic Ocean, paleoPacific
INTRODUCTION In North America, the history of the Appalachian-Ouachita orogen is commonly described in terms of the evolution of the Iapetus Ocean, the opening of which is recorded in the Late Neoproterozoic–Early Cambrian rifting of southeastern Laurentia (e.g., Cawood et al., 2001) and whose closure documents the accretion of a variety of peri-Gondwanan arc terranes (e.g., Ganderia, Avalonia, and Carolina; Murphy et al., 1995, 2004; van Staal et al., 1998; Hibbard, 2000) to the Laurentia margin in the Late Ordovician–Early Silurian (e.g., van Staal et al., 1998). Yet the climactic collision that produced the present-day Appalachian-Ouachita orogen was the result of the closure not of Iapetus, but of its successor, the Rheic Ocean. This ocean, which initially rifted in the Cambrian (e.g., Sánchez-García et al., 2003), is thought to have opened in the Early Ordovician with the separation of the peri-Gondwanan arc terranes from the margin of northern Gondwana documented by a period of rapid subsidence (e.g., Prigmore et al., 1997) and to have closed with the collision of this margin with Laurentia during the late Paleozoic assembly of Pangea (e.g., Martínez Catalán et al., 2002). That the role of the Rheic Ocean in the development of the Appalachian-Ouachita orogen has historically received less attention than that of its better-known forebear is largely a function of preservation. The orogen contains both the rifted margin and final suture of the Iapetus Ocean, so it preserves a complete record of its opening and closure. However, it preserves no such margin of the Rheic Ocean, the suture of which lies outboard of the accreted peri-Gondwanan terranes, so it either lies buried beneath the sediments of the Coastal Plain or was removed with the opening of the Atlantic Ocean and the Gulf of México. However, vestiges of the Rheic Ocean and its rifted continental margin may be preserved in the Acatlán and Oaxacan
Complexes of southern México (Fig. 1), which together make up the largest areas of Paleozoic and Precambrian geology in the country. The Acatlán Complex exposes Paleozoic rocks that represent remnants of an ocean (e.g., Ortega-Gutiérrez, 1975, 1978a), whereas the Oaxacan Complex exposes Mesoproterozoic (ca. 1 Ga) basement (e.g., Keppie et al., 2003a; Solari et al., 2003) overlain by early Paleozoic continental margin rocks of Gondwanan affinity (Robison and Pantoja-Alor, 1968). Potential correlatives of these two complexes also occur in northeastern México (Ortega-Gutiérrez, 1978b; Nance et al., this volume) in the form of the ophiolitic Granjeno Schist (Ramírez-Ramírez, 1992; Dowe et al., 2005) of the Sierra Madre terrane (Fig. 1A) and the ca. 1 Ga Novillo Gneiss (Cameron et al., 2004), the early Paleozoic cover of which is likewise of Gondwanan affinity (Boucot et al., 1997; Stewart et al., 1999). That all of these complexes preserve an important record of Paleozoic ocean opening and closure is not at issue, and their palinspastic significance has long been recognized (e.g., Yañez et al., 1991). But whether this ocean was Iapetus, the Rheic Ocean, or some other oceanic tract remains controversial. For example, Yañez et al. (1991) suggested that the Acatlán Complex provided a record of Laurentia-Gondwana collision in the Devonian, an event they equated to the Acadian belt of the Appalachian orogen. More recently, Ortega-Gutiérrez et al. (1999) reassigned this collisional event to the Late Ordovician–Early Silurian and suggested that the Acatlán Complex preserved a vestige of the Iapetus suture, a view first put forward by Ortega-Gutiérrez (1981) and subsequently developed by Talavera-Mendoza et al. (2005). But in a series of continental reconstructions published by Keppie and Ramos (1999) and Keppie (2004), the Acatlán Complex has been placed on the Gondwanan margin of the Rheic Ocean, the closure of which did not occur until the late Paleozoic. Discriminating between these mutually exclusive models is essential
Vestige of the Rheic Ocean in North America
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Figure 1. (A) Terrane map of México showing location of Acatlán Complex (after Keppie, 2004), and (B) geological map of the Acatlán Complex (modified from Ortega-Gutiérrez et al., 1999).
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for paleogeographic constructions and requires an understanding of the tectonothermal evolution of the Acatlán Complex. However, only recently, with the completion of detailed structural-kinematic studies (e.g., Elías-Herrera and Ortega-Gutiérrez, 2002; Malone et al., 2002) and the publication of new radiometric and faunal age constraints (e.g., Keppie et al., 2004a,b, 2006; Sánchez-Zavala et al., 2004; Talavera-Mendoza et al., 2005; Middleton et al., this volume; Miller et al., this volume), has such an understanding become possible. In this article we summarize the new data and present the case that the Acatlán Complex is, in fact, a remnant of the Rheic and paleo-Pacific Oceans and that it records a protracted history of rifting, magmatism, subduction, and collision associated with the amalgamation and breakup of Pangea. ACATLÁN COMPLEX The Acatlán Complex (Fig. 1B) comprises a repeatedly deformed assemblage of tectonically interleaved metasedimentary rocks, granitoid bodies, and metamorphosed mafic-ultramafic units that are traditionally considered to represent the Paleozoic basement of the Mixteca terrane (Campa and Coney, 1983). To the east the complex is faulted against Mesoproterozoic (ca. 1 Ga) granulite-facies gneisses of the Oaxacan Complex (e.g., OrtegaGutiérrez, 1981; Keppie et al., 2003a; Solari et al., 2003). These Grenville-age gneisses constitute the basement of the Oaxaquia terrane (Keppie, 2004) and are unconformably overlain by a latest Cambrian–Tremadocian passive margin assemblage of clastic and carbonate rocks that contain a Gondwanan faunal assemblage (Robison and Pantoja-Alor, 1968; Landing et al., 2006). The northsouth Caltepec fault zone that defines the boundary between these two terranes is a dextral transpressive structure that was active at ca. 276 Ma and is overstepped by the Early Permian (Leonardian) redbeds of the Matzitzi Formation (Elías-Herrera and OrtegaGutiérrez, 2002). To the south the Acatlán Complex is juxtaposed against Mesozoic and Tertiary metamorphic and plutonic rocks of the Xolapa Complex (Herrmann et al., 1994) along the east-west dextral Chacalapa fault, whereas to the west it is thrust over by Mesozoic island-arc rocks of the Guerrero terrane (Elías-Herrera et al., 2000) along the east-dipping Papalutla Thrust. The northern limit of the Mixteca terrane is obscured by Mesozoic–Cenozoic rocks and by the active Trans-Méxican volcanic belt. The geology of the Acatlán Complex (Fig. 1B) is based largely on the pioneering work of Ortega-Gutiérrez (1975, 1978a), who recognized within it two major tectonic assemblages, a predominantly metasedimentary Petlalcingo Subgroup and a predominantly igneous Acateco Subgroup, each of which was overlain by deformed metasedimentary rocks of the Tecomate, Patlanoaya, and Otates Formations. Although the Petlalcingo Subgroup has since been elevated to group status, the Acateco Subgroup has been renamed the Piaxtla Group (Ramírez-Espinosa, 2001), and, more recently, both groups have been designated suites (discussed later; Keppie et al., 2006), the threefold subdivision of the Acatlán Complex recognized by Ortega-Gutiérrez (1975, 1978a) remains in general use today.
As envisaged by Ortega-Gutiérrez et al. (1999), the two main tectonic assemblages of the Acatlán Complex are components of a deformed thrust nappe comprising (1) a parautochthonous lower plate (the Petlacingo Group) made up of a variety of repeatedly deformed, predominantly low-grade siliciclastic rocks, and (2) an allochthonous upper plate (the Piaxtla Group) composed of various high-grade assemblages that include locally eclogitic and blueschist-facies mafic-ultramafic rocks, amphibolite-facies metasedimentary units, and a suite of variably deformed megacrystic granitoid bodies (Fig. 2A). Following exhumation, both plates were overstepped by the uppermost unit of the Acatlán Complex, consisting of the volcanic-sedimentary Tecomate, Patlanoaya, and Otates Formations, the present structure of which (Fig. 2B) was attributed to subsequent infolding about major upright north-south folds. Following Ortega-Gutiérrez (1975, 1978a), the structurally lower Petlalcingo Group is subdivided into an upper Cosoltepec Formation, a middle Chazumba Formation, and a basal Magdalena Migmatite. The Chazumba Formation comprises polydeformed, amphibolite-facies metapsammite and metapelite that contain several mafic-ultramafic tectonic lenses. The structurally overlying Cosoltepec Formation, on the other hand, comprises greenschist-facies phyllite, quartzite, and minor mafic volcanic rocks. The Magdalena Migmatite is dominated by pelitic and psammitic lithologies with minor amphibolitic and calc-silicate units that have been metamorphosed in the amphibolite facies and pervasively migmatized. The Piaxtla Group is subdivided into the Xayacatlán Formation, a high-grade suite of mafic-ultramafic bodies, metasedimentary rocks, and migmatites, and the K-feldspar megacrystic Esperanza Granitoids with which they are tectonically interlayered (Fig. 2A). Although they do not constitute an ophiolite, eclogites and amphibolites of mid-ocean ridge basalt, ocean island basalt, and island-arc affinity within the Xayacatlán Formation indicate an important oceanic component (Meza-Figueroa, 1998; MezaFigueroa et al., 2003). The high-pressure metamorphism recorded in the Piaxtla Group was considered by Ortega-Gutiérrez et al. (1999) to have culminated in partial melting of the metasediments to produce the megacrystic Esperanza Granitoids. An imprecise Sm-Nd garnet–whole-rock age on eclogite from the Xayacatlán Formation (Yañez et al., 1991) was taken to suggest a Devonian age for this event. The overlying Tecomate, Patlanoaya, and Otates Formations consist of low-grade to unmetamorphosed slate, sandstone, conglomerate, and limestone, which Ortega-Gutiérrez (1978a) inferred to overlie the thrust nappe unconformably based on deformational grade and the occurrence of Esperanza-like granitoid pebbles in the conglomerates. A Devonian depositional age for the Tecomate Formation, originally assigned on the basis of poorly preserved faunas in the limestones, was supported by an imprecise lowerintercept U-Pb zircon age (Yañez et al., 1991) from the intrusive La Noria Granite (Fig. 2A). Latest Fammenian fossils have since been recovered from the Otates Formation (Vachard and Flores de Dios, 2002; Derycke-Katir et al., 2005), which lies unconformably
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Figure 2. (A) Simplified tectonostratigraphic column for the Acatlán Complex, and (B) east-west cross-section through Acatlán in the northern part of the complex, as envisioned by Ortega-Gutiérrez et al. (1999). Fm.—Formation.
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on units correlated with the Cosoltepec Formation. A structural contrast between the Cosoltepec Formation, which records three phases of penetrative deformation, and the Tecomate Formation, which records only two, has recently been documented by Malone et al. (2002). These authors also showed that both phases of deformation in the Tecomate Formation additionally affect the Totoltepec pluton, a calc-alkaline granitoid body from which concordant U-Pb zircon ages of 287 ± 2 Ma and 289 ± 1 Ma have been recorded (Yañez et al., 1991; Keppie et al., 2004b). An Early Permian age for this deformation and metamorphism is also suggested by a muscovite K-Ar age from the Tecomate Formation of 288 ± 14 Ma (Weber et al., 1997). The Acatlán Complex is unconformably overlain by the Leonardian Matzitzi Formation, which also oversteps the Oaxacan Complex. The Early Permian juxtaposition of these two complexes is further supported by a concordant U-Pb age of 276 ± 1 Ma from syntectonic migmatitic neosome developed along the bounding Caltepec fault zone (Elías-Herrera and Ortega-Gutiérrez, 2002). The Matzitzi Formation lacks significant metamorphism, but was locally strongly deformed before the deposition of Mesozoic strata (Hernández-Láscarez, 2004). Undeformed granitic dikes (the San Miguel dikes in Fig. 2A) dated at 175 ± 3 Ma and 172 ± 1 Ma (Rb/Sr and Sm/Nd, respectively; Yañez et al., 1991) locally cut the Chazumba Formation and the mafic-ultramafic lenses it contains. Unconformably overlying Middle Jurassic (Bajocian) rocks vary from continental to shallow marine shelf sediments and are common to both the Acatlán and the Oaxacan Complexes (e.g., Westermann et al., 1984; Sandoval and Westermann, 1986; SilvaPineda and Gonzalez-Gallardo, 1988). PREVIOUS TECTONIC INTERPRETATIONS OF THE ACATLÁN COMPLEX On the basis of broad tectonostratigraphic similarities to parts of the Appalachian-Ouachita orogen and its close proximity to the ca. 1 Ga basement of the Oaxacan Complex, the Acatlán Complex has traditionally been broadly interpreted in terms of Laurentia-Gondwana (Amazonia) collision (e.g., Ortega-Gutiérrez, 1981, 1993; Yañez et al., 1991). However, several conflicting models have been proposed to account for its tectonothermal history. Based on the imprecise Sm-Nd age for eclogites of the Piaxtla Group and the broadly Devonian U-Pb lower intercept age for the La Noria Granite (Fig. 2A), which was taken to be syntectonic and representative of the Esperanza Granitoids, Yañez et al. (1991) correlated the Acatlán Complex with the Acadian belt of the Appalachian orogen, suggesting that both were deformed during a Devonian collision (Mixtecan orogeny) between eastern Laurentia and northwestern South America (Sánchez-Zavala et al., 2000). The complex is thought to have then been transported southward with Gondwana to a position near present-day Colombia, recolliding with southern Laurentia during the late Carboniferous before moving to its present position following the breakup of Pangea. However, on the basis of a revised lower intercept age for the Esperanza Granitoids of 440 ± 14 Ma, Ortega-Gutiérrez et al.
(1999) proposed that the Acatlán Complex represented a vestige of the Iapetus suture (Fig. 3A) formed during a Late Ordovician– Early Silurian collision (Acatecan orogeny) between eastern Laurentia and Oaxaquia (Fig. 1A), a crustal block of Grenville age inferred to underlie much of present-day México (Ortega-Gutiérrez et al., 1995). Oaxaquia (of which the Oaxacan Complex is part) was considered to represent either a microcontinent or part of the Colombian margin of Gondwana. Emplacement of the allochthonous Piaxtla Group over the parautochthonous Petlacingo Group was attributed to this Late Ordovician–Early Silurian collision and interpreted to juxtapose eclogite-bearing oceanic rocks against trench and/or forearc deposits. The timing of this event also provided a minimum depositional age for the sedimentary rocks of both assemblages. The Late Devonian Mixtecan orogeny defined by the La Noria Granite, on the other hand, was held to be responsible for the development of the Magdalena Migmatite. The timing of this younger orogenic event was also thought to provide a minimum depositional age for the Tecomate Formation, because the granite was interpreted to intrude rocks of this formation syntectonically. Together with the age of the Totoltepec Granite (287 ± 2 Ma), the tectonothermal history of the complex showed striking parallels with that of the Appalachian orogen, lending support to its interpretation in terms of Appalachian tectonics and the closure of the Iapetus Ocean, a linkage recently expanded upon by Talavera-Mendoza et al. (2005). However, neither of these models is compatible with recent age data, which show much of the Acatlán Complex to be appreciably younger than previously thought. Instead, these new data support the proposition of Keppie and Ramos (1999) and Keppie (2004) that the Acatlán Complex represents a vestige not of Iapetus, but of its immediate successor, the Rheic Ocean (Fig. 3B), subsequently overprinted by convergent tectonics at the margin of the paleo-Pacific. If so, the complex may have lain adjacent to northwestern South America throughout the Paleozoic and its stratigraphic and tectonothermal history would provide an unique record of an important but often overlooked ocean whose closure produced the climactic Ouachita-Alleghanian-Variscan orogenic belt during the Permo-Carboniferous assembly of Pangea. NEW AGE CONSTRAINTS FOR THE ACATLÁN COMPLEX A variety of new age data, both paleontological and radiometric, have recently become available for the Acatlán Complex (Keppie et al., 2004a,b, 2006; Sánchez-Zavala et al., 2004; Talavera-Mendoza et al., 2005; Middleton et al., this volume; Miller et al., this volume). These data are summarized in Table 1. Depositional Age Constraints On the basis of existing age data, field relationships, and new LA-ICPMS (laser ablation inductively coupled plasma mass spectroscopy) detrital zircon U-Pb ages (Keppie et al., 2006), deposition of the various components of the Petlalcingo
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Group (redesignated a suite; Keppie et al., 2006), which were previously thought to predate emplacement of the Esperanza Granitoids at ca. 440 Ma (Ortega-Gutiérrez et al., 1999), has been constrained as follows: (1) deposition of the Cosoltepec Formation to the interval 455–360 Ma (Upper Ordovician–latest Devonian), although some parts may be no older than Lower Devonian (Talavera-Mendoza et al., 2005); (2) deposition of the Chazumba Formation (redesignated a lithodeme; Keppie et al., 2006) to the interval 239–173 Ma (Middle Triassic–Early Jurassic); and (3) deposition of the Magdalena Migmatite protolith (redesignated the Magdalena Lithodeme; Keppie et al., 2006) to the interval 303–171 Ma (Permian–Early Jurassic). Similarly, deposition of part of the Piaxtla Group (redesignated a suite; Keppie et al., 2006) has been constrained to the interval 700–470 Ma, while the presence of xenoliths of these rocks within the ca. 470 Ma megacrystic granitoids with which they are associated suggests an original intrusive relationship (Murphy et al., 2006). Newly discovered fossils and U-Pb ages from granitoid cobbles in the Tecomate Formation have constrained its depositional age, previously thought to predate emplacement of the La Noria Granite at ca. 371 Ma (Ortega-Gutiérrez et al., 1999), to the latest Pennsylvanian–earliest Middle Permian (Keppie et al., 2004a), although unfossiliferous rocks above and below the fossil-bearing limestones may extend its depositional age range. Detrital zircons from the Cosoltepec Formation (Keppie et al., 2006) show age clusters at ca. 470–455 Ma, ca. 660–500 Ma,
and ca. 1340–860 Ma and older ages of ca. 1735–1675 Ma, ca. 1880 Ma, ca. 1990 Ma, ca. 2155 Ma, and ca. 2725 Ma. The deposition of this formation is consequently constrained between the youngest concordant zircon, which has a 206Pb/238U age of 455 ± 4 Ma (Upper Ordovician; Gradstein et al., 2004), and the paleontologically defined sub-Strunian (latest Famennian) unconformity at the base of the Otates Formation (Vachard and Flores de Dios, 2002; Derycke-Katir et al., 2005). The ca. 410 Ma detrital zircon age cluster reported from the Cosoltepec Formation by Talavera-Mendoza et al. (2005), if representative, would constrain its deposition to the Devonian, but may show that the formation, as presently mapped, contains sedimentary packages of differing ages. Almost all of the detrital zircons from the Chazumba Lithodeme (Keppie et al., 2006) show ages in the range ca. 1150– 920 Ma, but the youngest has a concordant age of 239 ± 4 Ma (Middle Triassic; Gradstein et al., 2004). The youngest detrital zircon population reported from the Chazumba Formation by Talavera-Mendoza et al. (2005) peaks at ca. 275 Ma. Deposition of the formation is consequently constrained between these ages and that of the granitic San Miguel dikes (ca. 173 Ma; Yañez et al., 1991) that cut the unit. Zircons from the paleosome of the Magdalena Lithodeme show age clusters at ca. 930–840 Ma, ca. 1020–960 Ma, and ca. 1320–1060 Ma, with younger ages of ca. 300 Ma and ca. 520 Ma and older ages of ca. 1450 Ma and ca. 1575 Ma. Interpreted as detrital, these zircons constrain the deposition of the migmatite
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TABLE 1. SUMMARY OF CURRENT BEST-INTERPRETATION AGE CONSTRAINTS FOR ROCK UNITS IN THE ACATLÁN COMPLEX Lithotectonic unit/ Age Method/constraint deformational constraint (Ma) Tecomate Formation Youngest detrital zircon ≤300 – ≥270 Fossils Pennsylvanian–Middle Permian (earliest Guadelupian) Olinala Formation Youngest detrital zircon <286 Youngest detrital zircon peak <297 Piaxtla Suite Eclogite Xayacatlán Formation
La Noria Granite Los Hornos Granite Palo Liso Granite Esperanza Granitoids
Asis Lithodeme
Petlalcingo Suite Cosoltepec Formation
Chazumba Formation
Magdalena Proto lith
Magdalena migmatization Deformational constraints Formation of Tecomate (2 phases) Formation of Cosoltepec (3 phases) Exhumation of Piaxtla Suite (phase 1) Acatlán-Oaxaca juxtaposition (phase 2) Cosoltepec thrust over Chazumba (phase 3)
346 ± 3 442 ± 1 <447 <804 <700 >471 ca. 465 ca. 445 ca. 461 ± 2 ca. 440 ca. 461 ca. 471 ca. 478 ca. 460 ca. 478 ca. 471 >471 <700
<455 >Latest Devonian (ca. 365)<410 <341 >Pennsylvanian
U-Pb zircon U-Pb zircon Youngest detrital zircon Youngest detrital zircon Youngest detrital zircon U-Pb zircon on intrusive granitoid U-Pb zircon U-Pb zircon U-Pb zircon U-Pb zircon U-Pb zircon U-Pb zircon U-Pb zircon U-Pb zircon U-Pb zircon U-Pb zircon U-Pb zircon on intrusive granitoid Youngest detrital zircon
References 1 1 2 2
3 4 2 2 5 5 6 6 6 2 2 2 2 5 7 8 9 9
10 11, 12 2 2 1
<239 <249 >173 <303 <245 >171 171 ± 1
Youngest detrital zircon Fossils in overlying Otates Formation Youngest detrital zircon peak Youngest detrital zircon Fossils in overlying Tecomate Formation Youngest detrital zircon Youngest detrital zircon Cross-cutting San Miguel dikes Youngest detrital zircon Youngest detrital zircon U-Pb metamorphism U-Pb metamorphism
<287 ± 2 <289 ± 1 >Jurassic <410 >365 Jurassic <346 >365 276 ± 1
U-Pb zircon Totoltepec pluton U-Pb zircon Totoltepec pluton Undeformed cover Youngest detrital zircon Fossils in overlying Otates Formation Undeformed cover U-Pb zircon, eclogite Fossils in overlying sediments 40 39 Ar/ Ar Caltepec fault zone
13 14 15, 16 10 11, 12 15, 16 3 11 17
<239
10 2 13 14 2 14 14
Youngest detrital zircon in Chazumba 10 Formation. 40 39 224 ± 2 (?) Ar/ Ar muscovite from Cosoltepec 14 Formation References: 1—Keppie et al., 2004b; 2—Talavera-Mendoza et al., 2005; 3—Middleton et al., this volume; 4—Keppie et al., 2005; 5—Murphy et al., 2006; 6—Miller et al., this volume; 7—Campa-Uranga et al., 2002; 8—Sánchez-Zavala et al., 2004; 9—Murphy et al., 2005; 10—Keppie et al., 2006; 11—Vachard and Flores de Dios, 2002, 12—Derycke-Khatir et al., 2005; 13—Yañez et al., 1991; 14—Keppie et al., 2004b ; 15—Westermann et al., 1984; 16—Silva-Pineda and Gonzalez-Gallardo, 1988; 17—Elías-Herrera and Ortega-Gutiérrez, 2002.
Vestige of the Rheic Ocean in North America protolith to the interval between the youngest zircon, which has a concordant age of 303 ± 6 Ma (Upper Pennsylvanian; Gradstein et al., 2004), and the age of migmatization, which has recently been dated at 171 ± 1 Ma (Keppie et al., 2004b). This concurs with the detrital zircon ages reported from the Magdalena Lithodeme by Talavera-Mendoza et al. (2005), the youngest population of which peaks at ca. 317 Ma. Given the difference between the depositional age of the Cosoltepec Formation and those of the Chazumba and Magdalena Lithodemes, Keppie et al. (2006) recommend excluding the Cosoltepec Formation from the Petlalcingo Suite. Based on tectonic considerations, its distal turbiditic nature with continentalderived detritus, and, at least locally, the oceanic affinity of its mafic volcanics (e.g., Keppie et al., this volume), these authors propose a continental rise setting for the deposition of the Cosoltepec Formation along the southern (Gondwanan) margin of the Rheic Ocean (Fig. 4). By contrast, Keppie et al. (2006) suggest that the Chazumba and Magdalena Lithodemes, which make up the renamed Petlalcingo Suite, represent a turbiditic clastic wedge deposited in front of Permo-Triassic thrusts (discussed later) that emplaced the Totoltepec pluton and the Cosoltepec Formation onto the Tecomate and Chazumba Units, respectively. Murphy et al. (2006) report detrital zircon ages from a psammitic metasedimentary unit in part of the Piaxtla Suite (designated the Asis Lithodeme; Middleton et al., this volume), most of which cluster in the range ca. 1250–1050 Ma, but which also include a younger cluster at ca. 900 Ma and older ages of ca. 1330 Ma, ca. 1490 Ma, ca. 1550 Ma, and ca. 1660 Ma. However, the youngest zircon has a concordia age of 705 ± 8 Ma (Cryogenian; Gradstein et al., 2004). Deposition of these metasediments is consequently constrained between this age and that of a quartzaugen granite (ca. 470 Ma; concordant U-Pb SHRIMP [sensitive high-resolution microprobe] age) that contains xenoliths of the metasediments and is considered to intrude them. This matches the results of Talavera-Mendoza et al. (2005), who report zircons from the Xayacatlán Formation no younger than ca. 700 Ma. However, magmatic zircons reported by these authors from blueschist of the Piaxtla Suite show a peak at ca. 477 Ma, which is broadly contemporary with granitoid emplacement. The depositional age of part of the Tecomate Formation has recently been constrained by the discovery, in two separate limestone horizons, of the conodonts Sweetognathus subsymmetricus and Gondolella sp. and Neostreptognathodus sp. and Streptognathodus sp., respectively, the former indicating a latest Leonardian–earliest Guadelupian (ca. 276–268 Ma; Gradstein et al., 2004) depositional age, whereas the latter places the maximum age of deposition close to the Pennsylvanian-Permian boundary (Keppie et al., 2004a). Unfossiliferous rocks above and below the limestones may extend this depositional age range, but the paleontological constraints are supported by SHRIMP data for zircons from granite cobbles in a conglomerate of the Tecomate Formation, which have yielded concordant U-Pb ages as young as ca. 284 Ma, suggesting their derivation from the mid– Early Permian (287 ± 2 Ma) Totoltepec Granite.
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SIBERIA
Avalonia 550-750 Ma
Cadomia BALTICA
N Granjeno EA continental C O rise
Iberia
Florida basement ca. 550-625 Ma NORTH AMERICA
Yucatan basement ca. 540-560 Ma Amazon craton AFRICA
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EI
H
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Cosoltepec continental rise Chortis Oaxaquia ca. 920-1250 Ma
SOUTH AMERICA Arequipa Cuyania
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Chilenia
+
(ca. 420 Ma) Figure 4. Silurian (ca. 420 Ma) continental reconstruction of Keppie and Ramos (1999) showing the locations of the Acatlán and Oaxacan Complexes on the Gondwana margin of the Rheic Ocean (from Nance et al., this volume). The Granjeno continental rise refers to rocks of the Granjeno Schist in the Sierra Madre terrane of northeastern México with which the Cosoltepec Formation has been correlated (OrtegaGutiérrez, 1978b; Nance et al., this volume).
Tectonothermal Age Constraints On the basis of recent LA-ICMS, TIMS (thermal ionization mass spectrometry), and SHRIMP U-Pb zircon data, we have drawn the following conclusions: (1) the K-feldspar megacrystic granitoids, including the Esperanza Granitoids, previously assigned to the Late Ordovician–Early Silurian, and the those, like the La Noria Granite, previously thought to define a Late Devonian tectonothermal event (the Mixtecan orogeny of Ortega-Gutiérrez et al., 1999), have crystallization ages that span the Ordovician (Talavera-Mendoza et al., 2005; Miller et al., this volume); (2) the Xayacatlán Formation (Piaxtla Suite), in its type area, was deposited syntectonically in the earliest Silurian, although elsewhere (e.g., in the San Francisco de Asís area; Murphy et al., 2006) it predates megacrystic granitoid emplacement at ca. 470 Ma; (3) the high-grade metamorphism of the Piaxtla Suite (the Acatecan orogeny of Ortega-Gutiérrez et al., 1999) is, at least locally, Mississippian rather than Late Ordovician–Early Silurian in age (Middleton et al., this volume); and (4) the development of the Magdalena Migmatite, previously attributed to the Mixtecan orogeny, and the crystallization and metamorphism of mafic-ultramafic lenses in the Chazumba Formation, previously interpreted to be fragments of Iapetus (Ortega-Gutiérrez et al., 1999), record a thermal pulse of Jurassic age (Keppie et al., 2004b).
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Recent U-Pb zircon ages from the K-feldspar megacrystic Esperanza Granitoids of ca. 478 Ma (LA-ICPMS; CampaUranga et al., 2002), ca. 471 Ma (LA-ICMPS; Sánchez-Zavala et al., 2004), 480–470 Ma and 460–440 Ma (LA-ICPMS; Talavera-Mendoza et al., 2005), and ca. 460 Ma (SHRIMP; Miller et al., this volume) are similar, within error, to those obtained from the La Noria Granite (ca. 465 Ma, SHRIMP; Miller et al., this volume) and two other megacrystic granitoid bodies previously thought to intrude the Tecomate Formation, the Los Hornos Granite (ca. 445 Ma; Miller et al., this volume) and the Palo Liso Granite (461 ± 2 Ma, TIMS; Miller et al., this volume). These data suggest that, rather than representing syntectonic granitoid bodies of different (Late Ordovician–Early Silurian and Late Devonian) age emplaced during separate (Acatecan and Mixtecan) orogenic events, the K-feldspar megacrystic granitoid bodies form a single suite of plutons whose emplacement spans the Ordovician. Older (ca. 1165 Ma and 1043 Ma) ages reported by Talavera-Mendoza et al. (2005) are considered to date the Grenville-age inheritance that is present in all of the megacrystic granitoids. Geochemically, the granitoids are mainly of calc-alkaline affinity and straddle the boundary between the volcanic arc and within-plate fields on the discrimination diagrams of Pearce et al. (1984) (Reyes-Salas, 2003; Murphy et al., 2006; Miller et al., this volume). Keppie et al. (2005) report a concordant U-Pb zircon TIMS age of 442 ± 1 Ma and a virtually identical 40Ar/39Ar hornblende plateau age of 441 ± 15 Ma from an amphibolite of the Xayacatlán Formation (Piaxtla Suite) from the type area of this mafic unit at Xayacatlán. Together these data suggest that the crystallization of at least this member of the suite took place syntectonically in the earliest Silurian. Keppie et al. (2005) further show that this mafic unit has a continental tholeiitic affinity, suggesting its emplacement in a rifted arc or continental rift environment, and was metamorphosed under amphibolite-facies conditions, cooling though ~500 °C by ca. 441 Ma. Amphibolites from the Asís Lithodeme (part of the Piaxtla Suite) show continental tholeiitic affinities and are interlayered with continentally derived, immature rift-related to passive margin metasedimentary rocks, the deposition of which is considered to predate megacrystic granitoid emplacement at ca. 470 Ma (Murphy et al., 2006). Middleton et al. (this volume) report a concordant U-Pb zircon TIMS age of 346 ± 3 Ma from an eclogite in the Piaxtla Suite that has been confirmed by ca. 345 Ma SHRIMP analyses from a migmatite in the same suite that they infer to be the product of decompression melting. Together these ages show that the high-grade metamorphism of at least part of the Piaxtla Suite occurred in the Mississippian rather than the Late Ordovician–Early Silurian. Rapid exhumation is suggested by the fact that, 10 km north of the Asis area, latest Devonian rocks of the Otates Formation rest unconformably on the Piaxtla Suite (Elias-Herrera et al., 2004). The ages also require that this high-grade event (the Acatecan orogeny of Ortega-Gutiérrez et al., 1999) be entirely unrelated to the emplacement of the K-feldspar megacrystic granitoids.
Keppie et al. (2004b) report U-Pb zircon TIMS ages of 171 ± 1 Ma for the neosome of the Magdalena Migmatite and 174 ± 1 Ma for a metanorite from a mafic-ultramafic lens (the Tultitlan Amphibolite) in the Chazumba Formation. These data preclude the Paleozoic age previously inferred for these rocks (e.g., Ortega-Gutiérrez et al., 1999) and instead support earlier Rb-Sr and Sm-Nd ages of 175 ± 3 Ma and 172 ± 1 Ma, respectively, from the undeformed San Miguel Granite dikes that locally cut the Chazumba Formation (Yañez et al., 1991). Together they are taken to document an important Middle Jurassic thermal event that is interpreted to be plume-related and linked to the opening of the Gulf of México. Deformational Age Constraints Malone et al. (2002) have shown that, whereas three phases of penetrative deformation affected the Cosoltepec Formation, the Tecomate Formation was affected by only two, both of which are additionally recorded in the arc-related Totoltepec pluton, the felsic and mafic phases of which have yielded concordant U-Pb zircon TIMS ages of 287 ± 2 Ma (Yañez et al., 1991) and 289 ± 1 Ma (Keppie et al., 2004b), respectively. The oldest phase of deformation in the Cosoltepec Formation is bracketed between the youngest detrital zircon age from this formation (ca. 455 Ma; Keppie et al., 2006) and the latest Carboniferous maximum depositional age of the Tecomate Formation. This is consistent with the age of the youngest detrital zircon cluster (ca. 410 Ma) reported from the formation by Talvera-Mendoza et al., 2005). The Jurassic age of the postdeformational Mesozoic cover succession (Westermann et al., 1984; Silva-Pineda and GonzalezGallardo, 1988) constrains the second and third deformational phases to the Permo-Triassic. The first deformation produced a pervasive low-grade foliation axial planar to isoclinal folds and likely reflects the tectonic emplacement of the structurally overlying Piaxtla Suite, the exhumation of which is Late Devonian–Mississippian (Middleton et al., this volume). The second deformation produced a second low-grade foliation axial planar to isoclinal sheath folds and involved north-south dextral shearing and south-vergent thrusting during which the Totoltepec pluton was thrust over the Tecomate and Cosoltepec Formations. Identical deformational kinematics are developed along the Caltepec fault zone, where they record the tectonic juxtapositioning of the Acatlán and Oaxacan Complexes and have been dated as Early Permian (276 ± 1 Ma; ElíasHerrera and Ortega-Gutiérrez, 2002). The final deformational event produced a crenulation cleavage axial planar to broadly north-south-trending upright, open folds, the largest of which are responsible for the regional structure (Fig. 1). The event likely reflects the tectonic emplacement of the Cosoltepec Formation over the underlying Chazumba Lithodeme, the youngest detrital zircon from which (ca. 239 Ma; Keppie et al., 2006) constrains the event to the Middle–Late Triassic. This final event may be dated by a muscovite plateau age of 224 ± 2 Ma (early Upper Triassic; Gradstein et al., 2004) obtained from the Cosoltepec Formation
Vestige of the Rheic Ocean in North America near its tectonic contact with the Chazumba Lithodeme (Keppie et al., 2004b). REVISED TECTONIC INTERPRETATION OF THE ACATLÁN COMPLEX The new age constraints require significant modification of the tectonostratigraphy of the Acatlán Complex proposed by Ortega-Gutiérrez et al. (1999) (Fig. 2A). These revisions are shown schematically in Figure 5. Following Keppie et al. (2006), the Cosoltepec Formation has been separated from the Petlalcingo Suite, which comprises the Chazumba Lithodeme and the protolith of the Magdalena Migmatite (the Magdalena Lithodeme), and the main lithotectonic units have been positioned according to their approximate time of deposition and/or emplacement. Three major phases of thrust tectonics have been identified following the deposition of the rift-related to passive margin Piaxtla Suite in the Cambro-Ordovician and the continental rise deposits of the
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Cosoltepec Formation in the ?Siluro-Devonian: (1) Late Devonian–Mississippian juxtapositioning of the various components of the Piaxtla Suite and their emplacement over the Cosoltepec Formation prior to the deposition of the latest Devonian Otates Formation, the Permo-Carboniferous Patlanoaya Formation, and the latest Pennsylvanian–Middle Permian Tecomate Formation; (2) Early Permian exhumation and thrusting of the arc-related Totoltepec pluton over the Tecomate Formation during the tectonic juxtapositioning of the Acatlán and Oaxacan Complexes along the dextral Caltepec fault zone; and (3) Middle Triassic thrusting of the Tecomate and Cosoltepec Formations over the Chazumba and Magdalena Lithodemes. Based on these relationships, we interpret the Acatlán Complex to record the following Paleozoic tectonothermal history: (1) deposition of rift–passive margin sediments locally prior to ca. 470 Ma; (2) Ordovician–earliest Silurian continental arc–tholeiitic magmatism on the margin of Oaxaquia; (3) SiluroDevonian deposition of oceanic basalts and continental rise
100 Ortega-Gutiérrez et al. (1999)
This paper
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PETLALCINGO SUITE
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Patlanoaya Fm. (Late DevonianMiddle Permian)
Magdalena a migmatization mig
Mixtecan Orogeny (greenschist facies)
La Noria Granite (371 ± 34 Ma)
Tecomate Fm. (periarc) Acatecan Orogeny (eclogite facies)
PETLALCINGO GP.
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PALEO-PACIFIC Orogenic Event (370-340 Ma) (eclogite facies)
OAXACAN COMPLEX ca. 1 Ga
basalt
RHEIC OCEAN
Asis Lithodeme (Passive margin/ rift tholeiites)
A T
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Figure 5. Revised tectonostratigraphic column for the Acatlán Complex (Fm.—Formation; K—Cretaceous; J—Jurassic; Tr—Triassic; P—Permian; C—Carboniferous; D—Devonian; S—Silurian; O—Ordovician; –C—Cambrian; A—away; T—toward).
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deposits of the Cosoltepec Formation; (4) Late Devonian–early Carboniferous (≥360–345 Ma) subduction, eclogite-facies metamorphism, and migmatization of the Piaxtla Suite; (5) tectonic imbrication of the Piaxtla Suite and the Cosoltepec Formation followed by rapid exhumation of the Piaxtla Suite; (6) deposition and deformation of the latest Devonian–Middle Permian Otates, Patlanoaya, and Tecomate Formations before and during an Early Permian collisional episode that exhumed the arc-related Totoltepec pluton and juxtaposed the Acatlán Complex against Grenville-age basement of the Oaxacan Complex by way of dextral transpression and S-vergent thrusting; (7) Permo-Triassic deposition of the Chazumba-Magdalena clastic wedge in response to continued thrusting; and (8) Jurassic polyphase deformation, migmatization, metamorphism, and intrusion of mafic-ultramafic bodies related to passage over a plume (Keppie et al., 2004b). While the timing of the earlier tectonothermal events generally compares with those of the Appalachian-Ouachita orogen as earlier studies have claimed, the age data are clearly incompatible with a history linked to the Iapetus Ocean (e.g., Ortega- Gutiérrez et al., 1999), the Late Ordovician–Early Silurian closure of which predates that recorded within the Acatlán Complex. Instead we favor the reconstruction of Keppie and Ramos (1999) and Keppie (2004) and interpret the complex to be a vestige not of Iapetus, but of its immediate successor, the Rheic Ocean. Following the amalgamation of Pangea, the Acatlán Complex was subjected to tectonothermal events along the eastern margin of the paleo-Pacific Ocean. Prior to the opening of the Rheic Ocean, most recent palinspastic reconstructions place Avalonia adjacent to Oaxaquia and Amazonia on the northern margin of Gondwana (e.g., Keppie et al., 2003a, and references therein). Separation of Avalonia from this margin is generally thought to have occurred in the Early Ordovician (e.g., Cocks et al., 1997; Prigmore et al., 1997), although Landing (1996, 2004, 2005) has argued that Avalonia was microcontinental by the Early Cambrian on the basis of its distinct Avalonian fauna and the first episode of rapid subsidence recorded by its sedimentary rocks. Throughout the Cambrian, however, Avalonia remained peri-Gondwanan, as indicated by its Early Ordovician Gondwanan fauna. This fauna is gradually replaced by endemic forms in the Arenig–Llanvirn and by fauna of Baltic and Laurentian affinities in the Late Llanvirn–Ashgillian (Fortey and Cocks, 2003). Although many models suggest that the final separation of Avalonia was orthogonal (e.g., Dalziel, 1997; Stampfli and Borel, 2002), a model involving transtensional movement analogous to that of present-day Baja California has recently been proposed for its initial rifting by Keppie et al. (2003b). Murphy et al. (2006) and Miller et al. (this volume) therefore suggest that, following transtensional rifting and separation of Avalonia in the Cambrian, the Ordovician arc-like magmatism in the Acatlán Complex could record subduction along the southern margin of the Rheic Ocean. In this scenario, the early history of the Rheic Ocean overlapped that of Iapetus, although its Devono-Carboniferous closure provides a much closer match with the geologic history the Acatlán Complex records.
However, as these authors point out, the arc-related signature of the Ordovician megacrystic granitoids could equally well have been crustally inherited from the long history of subduction that the northern margin of Gondwana experienced in the Neoproterozoic (e.g., Murphy et al., 1999). If so, a more reliable indication of the tectonic setting may be recorded in the geochemistry of the mafic magmatism, which is more consistent with an extensional setting, possibly associated with Avalonia’s final separation from the peri-Gondwanan realm. Widespread megacrystic granitoid magmatism, for example, is associated with the Early Ordovician opening of the Rheic Ocean in the Central Iberian zone of Spain (e.g., Valverde-Vaquero and Dunning, 2000). Protracted rifting along the northern Gondwanan margin implied by the age span (ca. 480–445 Ma) of the megacrystic granitoids also parallels that of the European peri-Gondwanan margin (e.g., Stampfli et al., 2002) and would be consistent with models (e.g., Nance et al., 2002) that compare the Late Neoproterozoic–early Paleozoic history of this margin to the Mesozoic–Cenozoic history of the North American cordillera. In these models (1) the termination of subduction that characterized this margin in the Neoproterozoic is attributed to ridgetrench collision and the development of a continental transform analogous to the San Andreas system and (2) the initial rifting of peri-Gondwanan terranes such as West Avalonia and Carolina is thought to have occurred in an analogous fashion to the separation of Baja California from mainland México with the opening of the Sea of Cortez. This interpretation is consistent with the faunal data, which suggest that Avalonia had separated from Gondwana by the Early Cambrian (Landing, 1996, 2004, 2005). However, the faunal data also suggest that Avalonia remained a peri-Gondwanan microcontinent until the Early Ordovocian (Fortey and Cocks, 2003). If so, the separation of Avalonia from the northern Gondwanan margin may have occurred in two stages involving (1) initial rifting by the Early Cambrian, potentially followed by dispersal along the northern Gondwanan margin (e.g., Gutiérrez-Alonso et al., 2003), and (2) Early Ordovician drift that led to the opening of the Rheic Ocean. The margin-parallel movement of Avalonia during the Cambrian might thus be analogous to that of Baja British Columbia in the late Mesozoic–Cenozoic (e.g., Keppie and Dostal, 2001), whereas the eventual separation of Avalonia might have had the effect of removing a backstop, resulting in long-term extension of the thermally weakened and potentially overthickened Gondwanan crust in a fashion analogous to that of the Cenozoic Basin and Range province (e.g., Dickinson, 2002). Regardless of the tectonic setting of the Ordovician granitoids, the other components of the Piaxtla Suite probably represent vestiges of the Rheic Ocean and/or its marginal basins and the rifted margin of Oaxaquia, the Grenville-age basement of which has been linked to that of the Colombian Andes (Ruiz et al., 1999; Keppie et al., 2003a) and on which the oldest Paleozoic shallow marine strata are of latest Cambrian-Tremadocian age and Gondwanan affinity (Robison and Pantoja-Alor, 1968; Landing et al., 2006). Each of these features is consistent with deposition
Vestige of the Rheic Ocean in North America of the Cosoltepec Formation along the Gondwanan margin of the Rheic Ocean (Fig. 5), the leading edge of which (the Piaxtla Suite) was subducted and exhumed onto the Cosoltepec Formation during the Carboniferous amalgamation of Pangea (Fig. 6A). The Permo-Triassic events are probably related to convergent tectonics along the paleo-Pacific margin of Pangea. Subduction along this margin, which led to the development of a Permian arc (Torres et al., 1999), is recorded in the emplacement of the Totoltepec pluton (Malone et al., 2002). Accompanying oblique convergence is suggested by the rapid exhumation of this pluton, which supplied detritus to the broadly coeval Tecomate Formation (Fig. 6B). Subsequent Triassic shortening likely reflects continued convergence on this margin and was responsible for the exhumation of the Cosoltepec Formation (Fig. 6C) and the coeval deposition of the clastic wedge recorded in the rocks of the Chazumba and Magdalena Lithodemes (Keppie et al., 2006).
A
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PIAXTLA METAMORPHIC CX Rift and Rise
Totoltepec laccolith (arc) Upper part of MATZITZI & PATLANOAYA & TECOMATE IXTALTEPEC A T
DISCUSSION AND CONCLUSIONS
B Previous models for the evolution of the Acatlán Complex used the close match of available age data with tectonothermal events recorded in the Appalachian orogen to argue for a history linked to the closure of the Iapetus Ocean. Believing the earliest deformation to be of Devonian age, Yañez et al. (1991) interpreted the complex to record a collision (later termed the Mixtecan orogeny) between eastern Laurentia and northwestern Amazonia, which they considered equivalent to the Acadian orogeny. Subsequent identification of a ca. 440 Ma tectonothermal event led Ortega-Gutiérrez et al. (1999) to interpret the complex as a vestige of the Iapetus suture formed during Late Ordovician–Early Silurian collision (the Acatecan orogeny) between eastern Laurentia and Oaxaquia, a Grenville-age crustal block of Gondwanan affinity. A model involving the suturing of both peri-Laurentian and peri-Gondwanan arcs to Laurentia during Iapetus closure is likewise advocated by Talavera-Mendoza et al. (2005). As recorded in the Appalachian orogen, the Iapetus Ocean achieved rift-drift transition at the base of the Cambrian (e.g., Cawood et al., 2001), such that its Laurentian margin preserves shallow marine successions of Early Cambrian age (Thomas, 1991). Closure of Iapetus broadly coincided with the Acatecan orogeny of Ortega-Gutiérrez et al. (1999), but did not occur as the result of continent-continent collision between Laurentia and Gondwana. Instead, closure of Iapetus was the result, in the north, of a collision with Baltica and, in the south, of the accretion of smaller continental arc terranes (Neoproterozoic Ganderia, Avalonia, and Carolina) that had previously rifted away from their former positions along the margin of Gondwana (e.g., Murphy et al., 1995; van Staal et al., 1998; Hibbard, 2000; Keppie et al., 2003b). During terrane transfer, the portion of the Iapetus Ocean that lay ahead of these peri-Gondwanan terranes was closed, while a new Rheic Ocean opened behind them. Opening of the Rheic Ocean overlapped the closure of Iapetus and produced a new, rifted continental margin along the border of Gondwana. Subsidence curves for Avalonia have been taken to indicate an Early Ordovician age
EARLY PERMIAN
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S
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AT
PETLALCINGO SUITE Chazumba & Magdalena Lithodemes (clastic wedge)
Figure 6. Plate tectonic interpretation of the Acatlán Complex showing (A) Late Devonian–Mississippian subduction and exhumation of the Piaxtla Suite, an oceanic and rift-related continental margin assemblage (Tehuitzingo ophiolite from Proenza et al., 2004), onto the Gondwanan continental rise deposits of the Petlalcingo Group during closure of the Rheic Ocean between Oaxaquia and Laurentia; (B) Early Permian development of the Totoltepec arc and deposition and deformation of the Tecomate Formation during the collisional episode that juxtaposed the Acatlán and Oaxacan Complexes; and (C) Middle Triassic development of the Chazumba-Magdalena clastic wedge in response to continued subduction of the paleo-Pacific following the amalgamation of Pangea. (CX—Complex; FM—Formation; FZ— fault zone; S—south; A—away; T—toward).
for the Rheic rift-drift transition (Prigmore et al., 1997), although faunal data suggest a two-stage separation involving (1) initial rifting in the latest Neoproterozoic (Landing, 1996, 2004, 2005) and (2) final separation in the Arenig (Fortey and Cocks, 2003). Closure of the Rheic Ocean to form the Alleghanian-Ouachita orogen was the result of collision in the late Paleozoic between Laurentia (and its accreted terranes) and Gondwana (e.g., Hatcher, 2002) during the amalgamation of Pangea. On the basis of the new age data, the high-grade (“Acatecan”) metamorphism of the Piaxtla Suite has been shown to be
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Late Devonian–Mississippian, whereas the older (Ordovician– Early Silurian) ages from the Esperanza Granitoids and the Xayacatlán Formation have been taken to date the formation either of a continental rift- or arc-related assemblage or of a Basin and Range–style extensional regime along the northern Gondwanan margin. Terminal collision in the Devono-Carboniferous is consistent with the appearance of midcontinental (U.S.) fauna in the Mississippian rocks resting on Oaxaquia (Stewart et al., 1999; Navarro-Santillan et al., 2002). This history of late Paleozoic subduction and collision conflicts with that of Iapetus, but closely matches that of the Rheic Ocean. Rifting from the northern margin of Gondwana of the Hun terrane of southern Europe, which was specifically extended to include the peri-Gondwanan elements of Middle America by Stampfli et al. (2002), postdates that recorded in the Acatlán Complex and instead records the Late Silurian opening of an ocean (Paleotethys) whose closure did not occur until the Pennsylvanian (Stampfli and Borel, 2002). As a vestige of the Rheic Ocean, the Acatlán Complex can be interpreted to record events along the leading edge of Gondwana on the ocean’s southern margin during the amalgamation of Pangea. In this view, the high-grade (“Acatecan”) metamorphism of the Piaxtla Suite and its emplacement onto the lowgrade Cosoltepec Formation at ca. 360–345 Ma documents the subduction and exhumation of the leading edge of Gondwana concurrent with the closure of the Rheic Ocean. After the amalgamation of Pangea and the tectonic juxtapositioning of the Acatlán and Oaxacan Complexes, the development of a Permian magmatic arc on the southwestern (according to the present coordinates) flank of Pangea is recorded in the Acatlán Complex by the emplacement of the Totoltepec Granite and the deposition of the arc-related Tecomate Formation. This arc is considered to have developed in response to subduction of the paleo-Pacific following the assembly of Pangea. Continued convergent tectonics along this margin during the Permian and the Triassic likely led to the tectonic juxtapositioning of the Cosoltepec and Tecomate Formations and the development of the Chazumba-Magdalena clastic wedge and was followed in the Jurassic by hotspot activity associated with the overriding of a mantle plume during the opening of the Gulf of México. ACKNOWLEDGMENTS Funding for this project was provided by National Science Foundation grants EAR-0308105 to RDN and EAR-0308437 to BVM, PAPIIT (Programa de Apoyo a Proyectos de Investigación e Innovación Tecnológica) grant IN103003 and CONACyT (Consejo Nacional de Ciencia y Tecnología) funding to JDK, and NSERC (Natural Sciences and Engineering Research Council of Canada) Discovery Grants to JBM and JD. For introducing us to the Acatlán Complex and for continued discussions and support we are especially indebted to Dr. Fernando Ortega-Gutiérrez. Thoughtful reviews by Jürgen von Raumer and John Winchester greatly improved this article, which is a contribution to IGCP (International Geoscience Program) Projects 453 and 497.
REFERENCES CITED Boucot, A.J., Blodgett, R.B., and Stewart, J.H., 1997, European Province Late Silurian brachiopods from the Ciudad Victoria area, Tamaulipas, northeastern Mexico, in Klapper, G., Murphy, M.A., and Talent, J.A. eds., Paleozoic sequence stratigraphy, biostratigraphy, and biogeography: Studies in honor of J. Granville (“Jess”) Johnson: Boulder, Colorado, Geological Society of America Special Paper 321, p. 273–293. Cameron, K.L., Lopez, R., Ortega-Gutiérrez, F., Solari, L.A., Keppie, J.D., and Schulze, C., 2004, U-Pb constraints and Pb isotopic compositions of leached feldspars: Constraints on the origin and evolution of Grenvillian rocks from eastern and southern Mexico, in Tollo, R.P., et al., eds., Proterozoic tectonic evolution of the Grenville orogen in North America: Boulder, Colorado, Geological Society of America Memoir 197, p. 755–770. Campa, M.F., and Coney, P.J., 1983, Tectono-stratigraphic terranes and mineral distributions in Mexico: Canadian Journal of Earth Sciences, v. 20, p. 1040–1051. Campa-Uranga, M.F., Gehrels, G., and Torres de León, R., 2002, Nuevas edades de granitoides metamorfizados del complejo Acatlán en el estado de Guerrero: Actas Instituto Nacional de Geoquímica, v. 8, no. 1, p. 248. Cawood, P.A., McCausland, P.J.A., and Dunning, G.R., 2001, Opening Iapetus: Constraints from the Laurentian margin of Newfoundland: Geological Society of America Bulletin, v. 113, p. 443–453, doi: 10.1130/00167606(2001)113<0443:OICFTL>2.0.CO;2. Cocks, L.R.M., McKerrow, W.S., and van Staal, C.R., 1997, The margins of Avalonia: Geological Magazine, v. 134, p. 627–636, doi: 10.1017/ S0016756897007425. Dalziel, I., 1997, Overview: Neoproterozoic–Paleozoic geography and tectonics: Review, hypothesis, environmental speculation: Geological Society of America Bulletin, v. 109, p. 16–42, doi: 10.1130/0016-7606(1997)109<0016: ONPGAT>2.3.CO;2. Derycke-Katir, C., Vachard, D., Dégardin, J.-M., Flores de Dios, A., Buitrón, B., and Hansen, M., 2005, Late Pennsylvanian and Early Permian chondichthyan microremains from San Salvador Patlanoaya (Puebla, Mexico): Geobios, v. 38, p. 43–55, doi: 10.1016/j.geobios.2003.06.008. Dickinson, W.R., 2002, The Basin and Range Province as a composite extensional domain: International Geology Review, v. 44, p. 1–38. Dowe, D.S., Nance, R.D., Keppie, J.D., Cameron, K.L., Ortega-Rivera, A., Ortega-Gutiérrez, F., and Lee, J.W.K., 2005, Deformational history of the Granjeno Schist, Ciudad Victoria, Mexico: Constraints on the closure of the Rheic Ocean?: International Geology Review, v. 47, p. 920–937. Elías-Herrera, M., and Ortega-Gutiérrez, F., 2002, Caltepec fault zone: An Early Permian dextral transpressional boundary between the Proterozoic Oaxacan and Paleozoic Acatlán complexes, southern Mexico, and regional implications: Tectonics, v. 21, no. 3, doi: 10.1029/200TC001278, doi: 10.1029/2000TC001278. Elías-Herrera, M., Sánchez-Zavala, J.L., and Macías-Romo, C., 2000, Geologic and geochronologic data from the Guerrero terrane in the Tejupilco area, southern Mexico: New constraints on its tectonic interpretation: Journal of South American Earth Sciences, v. 13, p. 355–376, doi: 10.1016/S08959811(00)00029-8. Elias-Herrera, M., Ortega-Gutiérrez, F., Sanchez-Zavala, J.L., Reyes-Salas, A.M., Macias-Romo, C., and Alexander, I., 2004, New geochronological and stratigraphic data related to the Paleozoic evolution of the highpressure Piaxtla Group, Acatlán Complex, southern Mexico: IV Reunión Nacional de Ciencias de la Tierra, Libro de Resúmes, p. 150 (abstract). Fortey, R.A., and Cocks, L.R.M., 2003, Paleontological evidence bearing on global Ordovician–Siluran continental reconstructions: Earth-Science Reviews, v. 61, p. 245–307, doi: 10.1016/S0012-8252(02)00115-0. Gradstein, F.M., Ogg, J.G., Smith, A.G., Bleeker, W., and Lourens, L.J., 2004, A new Geologic Time Scale, with special reference to Precambrian and Neogene: Episodes, v. 27, p. 83–100. Gutiérrez-Alonso, G., Fernández-Suárez, J., Jeffries, T.E., Jenner, G.A., Tubrett, M.N., Cox, R., and Jackson, S.E., 2003, Terrane accretion and dispersal in the northern Gondwana margin. An Early Paleozoic analogue of a long-lived active margin: Tectonophysics, v. 365, p. 221–232, doi: 10.1016/S0040-1951(03)00023-4. Hatcher, R.D., Jr., 2002, Alleghanian (Appalachian) orogeny, a product of zipper tectonics: Rotational transpressive continent-continent collision and closing of ancient oceans along irregular margins, in Martínez Catalán, J.R., et al., eds., Variscan-Appalachian dynamics: The building of the late
Vestige of the Rheic Ocean in North America Paleozoic basement: Boulder, Colorado, Geological Society of America Special Paper 364, p. 199–208. Hernández Láscarez, D., 2004, Geología de la region de Los Reyes Metzontal– Santiago Coatepec, extremo suroriente del estado de Puebla, México: IV Reunión Nacional de Ciencias de la Tierra, Libro de Resúmes, p. 95 (abstract). Herrmann, U.R., Nelson, B.K., and Ratschbacher, L., 1994, The origin of a terrane: U/Pb geochronology and tectonic evolution of the Xolapa complex (southern Mexico): Tectonics, v. 13, p. 455–474, doi: 10.1029/93TC02465. Hibbard, J., 2000, Docking Carolina: Mid-Paleozoic accretion in the southern Appalachians: Geology, v. 28, p. 127–130, doi: 10.1130/0091-7613(2000) 028<0127:DCMPAI>2.3.CO;2. Keppie, J.D., 2004, Mexican terranes revisited: A 1.3 Ga odyssey: International Geology Review, v. 46, p. 765–794. Keppie, J.D., and Dostal, J., 2001, Evaluation of the Baja controversy using paleomagnetic and faunal data, plume magmatism, and piercing points: Tectonophysics, v. 339, p. 427–442, doi: 10.1016/S0040-1951(01)00126-3. Keppie, J.D., and Ramos, V.S., 1999, Odyssey of terranes in the Iapetus and Rheic Oceans during the Paleozoic, in Ramos, V.A., and Keppie, J.D., eds., Laurentian-Gondwanan connection before Pangea: Boulder, Colorado, Geological Society of America Special Paper 336, p. 267–276. Keppie, J.D., Dostal, J., Cameron, K.L., Solari, L.A., Ortega-Gutiérrez, F., and Lopez, R., 2003a, Geochronology and geochemistry of Grenvillian igneous suites in the northern Oaxaca Complex, southern Mexico—Tectonic implications: Precambrian Research, v. 120, p. 365–389, doi: 10.1016/ S0301-9268(02)00166-3. Keppie, J.D., Nance, R.D., Murphy, J.B., and Dostal, J., 2003b, Tethyan, Mediterranean, and Pacific analogues for the Neoproterozoic–Paleozoic birth and development of peri-Gondwanan terranes and their transfer to Laurentia and Laurussia: Tectonophysics, v. 365, nos. 1–4, p. 195–219, doi: 10.1016/S0040-1951(03)00037-4. Keppie, J.D., Sandberg, C.A., Miller, B.V., Sánchez-Zavala, J.L., Nance, R.D., and Poole, F.G., 2004a, Implications of latest Pennsylvanian to Middle Permian paleontological and U-Pb SHRIMP data from the Tecomate Formation to re-dating tectonothermal events in the Acatlán Complex, southern Mexico: International Geology Review, v. 46, p. 745–753. Keppie, J.D., Nance, R.D., Powell, J.T., Mumma, S.A., Dostal, J., Fox, D., Muise, J., Ortega-Rivera, A., Miller, B.V., and Lee, J.W.K., 2004b, MidJurassic tectonothermal event superposed on a Paleozoic geological record in the Acatlán Complex of southern Mexico: Hotspot activity during the breakup of Pangea: Gondwana Research, v. 7, p. 239–260, doi: 10.1016/S1342-937X(05)70323-3. Keppie, J.D., Dostal, J., Miller, B.V., Nance, R.D., Murphy, J.B., and Cooper, P. 2005, Geochemistry and geochronology of the Xayacatlán type area, Acatlán Complex, southern Mexico: ~440 Ma syntectonic emplacement of a layered continental tholeiitic gabbro during the birth of the Rheic Ocean: Geological Association of Canada Abstracts, v. 30, p. 101. Keppie, J.D., Nance, R.D., Fernández-Suárez, J., Storey, C.D., Jeffries, T.E., and Murphy, J.B., 2006, Detrital zircon data from the eastern Mixteca terrane, southern Mexico: Evidence for an Ordovician–Mississippian continental rise and a Permo-Triassic clastic wedge adjacent to Oaxaquia: International Geology Review, v. 48, p. 97–111. Keppie, J.D., Dostal, J., and Eliás-Herrera, M., 2007 (this volume), Ordovician– Devonian oceanic basalts in the Cosoltepec Formation, Acatlán Complex, southern México: Vestiges of the Rheic Ocean? in Linnemann, U., et al., eds., The evolution of the Rheic Ocean: From Avalonian-Cadomian active margin to Alleghenian-Variscan collision: Boulder, Colorado, Geological Society of America Special Paper 423, doi: 10.1130/2007.2423(21). Landing, E., 1996, Avalon: Insular continent by the latest Precambrian, 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. 29–63. Landing, E., 2004, Precambrian–Cambrian boundary interval deposition and the marginal platform of the Avalon microcontinent: Journal of Geodynamics, v. 37, p. 411–435, doi: 10.1016/j.jog.2004.02.014. Landing, E., 2005, Early Paleozoic Avalon-Gondwana unity: An obituary—Response to “Palaeontological evidence bearing on global Ordovician-Silurian continental reconstructions” by R.A. Fortey and L.R.M. Cocks: Earth-Science Reviews, v. 69, p. 169–175, doi: 10.1016/ j.earscirev.2004.10.002. Landing, E., Keppie, J.D., and Westrop, S.R., 2006, Lower Paleozoic of northwest Gondwana: Terminal Cambrian and lowest Ordovician Tiñu
451
Formation in southern Mexico: Geological Society of America, Abstracts with Programs, v. 38, no. 2, p. 21. Malone, J.W., Nance, R.D., Keppie, J.D., and Dostal, J., 2002, Deformational history of part of the Acatlán Complex: Late Ordovician–Early Silurian and Early Permian orogenesis in southern Mexico: Journal of South American Earth Sciences, v. 15, p. 511–524, doi: 10.1016/S08959811(02)00080-9. Martínez Catalán, J.R., Hatcher, R.D., Jr., Arenas, R., and Díaz García, F., eds., 2002, Variscan-Appalachian dynamics: The building of the late Paleozoic basement: Boulder, Colorado, Geological Society of America Special Paper 364, 312 p. Meza-Figueroa, D.M., 1998, Geochemistry and characterization of intermediate temperature eclogites from the Acatlan Complex, southern Mexico [Ph.D. thesis]: Tucson, University of Arizona, 201 p. Meza-Figueroa, D., Ruiz, J., Talavera-Mendoza, O., and Ortega-Gutierrez, F., 2003, Tectonometamorphic evolution of the Acatlan Complex eclogites (southern Mexico): Canadian Journal of Earth Sciences, v. 40, p. 27–44, doi: 10.1139/e02-093. Middleton, M., Keppie, J.D., Murphy, J.B., Miller, B.V., Nance, R.D., OrtegaRivera, A., and Lee, J.K.W., 2007 (this volume), P-T-t constraints on exhumation following subduction in the Rheic Ocean from eclogitic rocks in the Acatlan Complex of southern Mexico, in Linnemann, U., et al., eds., The evolution of the Rheic Ocean: From AvalonianCadomian active margin to Alleghenian-Variscan collision: Boulder, Colorado, Geological Society of America Special Paper 423, doi: 10.1130/2007.2423(21). Miller; B.V., Dostal, J., Keppie, J.D., Nance, R.D., Ortega-Rivera, A., and Lee, J.W.K., 2007 (this volume), Ordovician calc-alkaline granitoids in the Acatlán Complex, southern Mexico: Geochemical and geochronological data and implications for the tectonics of the Gondwanan margin of the Rheic Ocean, in Linnemann, U., et al., eds., The evolution of the Rheic Ocean: From Avalonian-Cadomian active margin to Alleghenian-Variscan collision: Boulder, Colorado, Geological Society of America Special Paper 423, doi: 10.1130/2007.2423(21). Murphy, J.B., Nance, R.D., Keppie, J.D., Dostal, J., and Cousens, B.L., 1995, Odyssey of West Avalonia: Isotopic constraints for late Proterozoic III–early Silurian paleogeography: St. John’s, Geological Association of Canada Special Paper 41, p. 227–238. Murphy, J.B., Keppie, J.D., Dostal, J., and Nance, R.D., 1999, Neoproterozoic– early Paleozoic evolution of Avalonia, in Ramos, V.A., and Keppie, J.D., eds., Laurentia-Gondwana connections before Pangea: Boulder, Colorado, Geological Society of America Special Paper 336, p. 253–266. Murphy, J.B., Pisarevsky, S.A., Nance, R.D., and Keppie, J.D., 2004, Neoproterozoic–Early Paleozoic configuration of peri-Gondwanan terranes: Implications for Laurentia-Gondwana connections: International Journal of Earth Sciences, v. 93, p. 659–682, doi: 10.1007/s00531-004-0412-9. Murphy, J.B., Keppie, J.D., Nance, R.D., Miller, B.V., Middleton, M., Fernández-Suarez, J., and Jeffries, T.E., 2006, Geochemistry and U-Pb protolith ages of eclogitic rocks of the Asís Lithodeme, Piaxtla Suite, Acatlán Complex, southern Mexico: Tectonothermal activity along the southern margin of the Rheic Ocean: Journal of the Geological Society of London, v. 163, p. 683–695. Nance, R.D., Murphy, J.B., and Keppie, J.D., 2002, A Cordilleran model for the evolution of Avalonia: Tectonophysics, v. 352, p. 11–31, doi: 10.1016/ S0040-1951(02)00187-7. Nance, R.D., Fernández-Suárez, J., Keppie, J.D., Storey, C., and Jeffries, T.E., 2007 (this volume), Provenance of the Granjeno Schist, Ciudad Victoria, México: Detrital zircon U-Pb age constraints and implications for the Paleozoic paleogeography of the Rheic Ocean, in Linnemann, U., et al., eds., The evolution of the Rheic Ocean: From Avalonian-Cadomian active margin to Alleghenian-Variscan collision: Boulder, Colorado, Geological Society of America Special Paper 423, doi: 10.1130/2007.2423(21). Navarro-Santillan, D., Sour-Tovar, F., and Centeno-Garcia, E., 2002, Lower Mississippian (Osagean) brachiopods from the Santiago Formation, Oaxaca, Mexico: Stratigraphic and tectonic implications: Journal of South American Earth Sciences, v. 15, p. 327–336, doi: 10.1016/S08959811(02)00047-0. Ortega-Gutiérrez, F., 1975, The pre-Mesozoic geology of the Acatlán area, south Mexico [Ph.D. thesis]: Leeds, England, University of Leeds, 166 p. Ortega-Gutiérrez, F., 1978a, Estratigrafia del Complejo Acatlán en la Mixteca Baja, Estados de Puebla y Oaxaca: Universidad Nacional Autónoma de México, Instituto de Geología, Revista, v. 2, p. 112–131.
452
Nance et al.
Ortega-Gutiérrez, F., 1978b, El Gneiss Novillo y rocas metamorficas asociadas en los canones del Novillo y la Peregrina, area Ciudad Victoria, Tamaulipas: Universidad Nacional Autónoma de México, Instituto de Geología, Revista, v. 2, p. 19–30. Ortega-Gutiérrez, F., 1981, Metamorphic belts of southern Mexico and their tectonic significance: Geofisica Internacional, v. 20, p. 177–202. Ortega-Gutiérrez, F., 1993, Tectonostratigraphic analysis and significance of the Paleozoic Acatlán Complex of southern México, in Ortega-Gutiérrez, F., et al., eds., First Circum-Pacific and Circum-Atlantic Terrane Conference, Guanajuato, México: México, D.F., Universidad Nacional Autónoma de México, Instituto de Geología, p. 54–60. Ortega-Gutiérrez, F., Ruiz, J., and Centeno-Garcia, E., 1995, Oaxaquia, a Proterozoic microcontinent accreted to North America during the late Paleozoic: Geology, v. 23, p. 1127–1130, doi: 10.1130/0091-7613(1995)023<1127: OAPMAT>2.3.CO;2. Ortega-Gutiérrez, F., Elias-Herrera, M., Reyes-Salas, M., Macias-Romo, C., and López, R., 1999, Late Ordovician–Early Silurian continental collision orogeny in southern Mexico and its bearing on Gondwana-Laurentia connections: Geology, v. 27, p. 719–722, doi: 10.1130/00917613(1999)027<0719:LOESCC>2.3.CO;2. Prigmore, J.K., Butler, A.J., and Woodcock, N.H., 1997, Rifting during separation of Eastern Avalonia from Gondwana: Evidence from subsidence analysis: Geology, v. 25, p. 203–206, doi: 10.1130/0091-7613(1997)025<0203: RDSOEA>2.3.CO;2. Proenza, J.A., Ortega-Gutiérrez, F., Camprubí, A., Tritlla, J., Elías-Herrera, M., and Reyes-Salas, M., 2004, Paleozoic serpentinite-enclosed chromitites from Tehuitzingo (Acatlán Complex, southern Mexico): A petrological and mineralogical study: Journal of South American Earth Sciences, v. 16, p. 649–666, doi: 10.1016/j.jsames.2003.12.003. Ramírez-Espinosa, J., 2001, Tectono-magmatic evolution of the Paleozoic Acatlan Complex in southern Mexico, and its correlation with the Appalachian system [Ph.D. thesis]: Tucson, University of Arizona, 170 p. Ramírez-Ramírez, C., 1992, Pre-Mesozoic geology of Huizachal-Peregrina Anticlinorium, Ciudad Victoria, Tamaulipas, and adjacent parts of eastern Mexico [Ph.D. thesis]: Austin, University of Texas, 317 p. Reyes-Salas, A.M., 2003, Mineralogia y petrologia de los granitoides Esperanza del compljo Acatlán, sur de México [Ph.D. thesis]: México City, Universidad Nacional Autónoma de México, 165 p. Robison, R., and Pantoja-Alor, J., 1968, Tremadocian trilobites from Nochixtlan region, Oaxaca, Mexico: Journal of Paleontology, v. 42, p. 767–800. Ruiz, J., Tosdal, R.M., Restrepo, P.A., and Murillo-Muneton, G., 1999, Pb isotope evidence for Columbia–southern Mexico connections in the Proterozoic, in Ramos, V.A., and Keppie, J.D., eds., Laurentian-Gondwanan connection before Pangea: Boulder, Colorado, Geological Society of America Special Paper 336, p. 183–197. Sánchez-García, T., Bellido, F., and Quesada, C., 2003, Geodynamic setting and geochemical signatures of Cambrian–Ordovician rift-related igneous rocks (Ossa-Morena Zone, SW Iberia): Tectonophysics, v. 365, p. 233– 255, doi: 10.1016/S0040-1951(03)00024-6. Sánchez-Zavala, J.L., Ortega-Gutiérrez, F., and Elías-Herrera, M., 2000, La orogenia Mixteca del Devónico del complejo Acatlán, sur de Mexico: GEOS, Unión Geofísica Mexicana, Boletin Informativo Época II, v. 20, no. 3, p. 321–322. Sánchez-Zavala, J.L., Ortega-Gutiérrez, F., Keppie, J.D., Jenner, G.A., Belousova, E., and Maciás-Romo, C., 2004, Ordovician and Mesoproterozoic zircons from the Tecomate Formation and Esperanza granitoids, Acatlán Complex, southern Mexico: Local provenance in the Acatlán and Oaxacan complexes: International Geology Review, v. 46, p. 1005–1021. Sandoval, J., and Westermann, G.E.G., 1986, The Bajocian (Jurassic) ammonite fauna of Oaxaca, Mexico: Journal of Paleontology, v. 60, p. 1220–1271. Silva-Pineda, A., and Gonzalez-Gallardo, S., 1988, Algunas Bennettitales (Cycadophyta) y Coniferales (Conifrophyta) del Jurasico medio del area de Cualac, Guerrero: Revista, Instituto de Geología, v. 7, no. 2, p. 244–248.
Solari, L.A., Keppie, J.D., Ortega-Gutiérrez, F., Cameron, K.L., Lopez, R., and Hames, W.E., 2003, 990 and 1100 Ma Grenvillian tectonothermal events in the northern Oaxacan Complex, southern Mexico: Roots of an orogen: Tectonophysics, v. 365, p. 257–282, doi: 10.1016/S0040-1951(03)00025-8. Stampfli, G.M., and Borel, G.D., 2002, A plate tectonic model for the Paleozoic and Mesozoic constrained by dynamic plate boundaries and restored synthetic oceanic isochrons: Earth and Planetary Science Letters, v. 196, p. 17–33, doi: 10.1016/S0012-821X(01)00588-X. Stampfli, G.M., von Raumer, J., and Borel, G.D., 2002, Palaeozoic evolution of pre-Variscan terranes: From peri-Gondwana to the Variscan collision, in Martinez-Catalan, J.R., et al., eds., Variscan-Appalachian dynamics: The building of the Upper Paleozoic basement: Boulder, Colorado, Geological Society of America Special Paper 364, p. 263–280. Stewart, J.H., Blodgett, R.B., Boucot, A.J., Carter, J.L., and Lopez, R., 1999, Exotic Paleozoic strata of Gondwanan provenance near Ciudad Victoria, Tamaulipas, Mexico, in Ramos, V.A., and Keppie, J.D., eds., LaurentiaGondwana connections before Pangea: Boulder, Colorado, Geological Society of America Special Paper 336, p. 227–252. Talavera-Mendoza, O., Ruíz, J., Gehrels, G.E., Meza-Figueroa, D.M., VegaGranillo, R., and Campa-Uranga, M.F., 2005, U-Pb geochronology of and implications for the Paleozoic paleogeography and tectonic evolution of southern Mexico: Earth and Planetary Science Letters, v. 235, p. 682–699, doi: 10.1016/j.epsl.2005.04.013. Thomas, W.A., 1991, The Appalachian-Ouachita rifted margin of southeastern North America: Geological Society of America Bulletin, v. 103, p. 415– 431, doi: 10.1130/0016-7606(1991)103<0415:TAORMO>2.3.CO;2. Torres, R., Ruíz, J., Patchett, P.J., and Grajales-Nishimura, J.M., 1999, PermoTriassic continental arc in eastern Mexico: Tectonic implications for reconstructions of southern North America, in Bartolini, C., Wilson, J.L., and Lawton, T.F., eds., Mesozoic sedimentary and tectonic history of north-central Mexico: Boulder, Colorado, Geological Society of America Special Paper 340, p. 191–196. Vachard, D., and Flores de Dios, A., 2002, Discovery of latest Devonian / earliest Mississippian microfossils in San Salvador Patlanoaya (Puebla, Mexico): Biogeographic and geodynamic consequences: Comptes Rendus Geoscience, v. 334, p. 1095–1101, doi: 10.1016/S1631-0713(02)01851-5. Valverde-Vaquero, P., and Dunning, G.R., 2000, New U-Pb ages from Early Ordovician magmatism in Central Spain: Journal of the Geological Society of London, v. 157, p. 15–26. van Staal, C.R., Dewey, J.F., Mac Niocaill, C., and McKerrow, W.S., 1998, The Cambrian–Silurian tectonic evolution of the Northern Appalachians and British Caledonides: History of a complex, west and southwest Pacifictype segment of Iapetus, in Blundell, D., and Scott, A.C., eds., Lyell: The past is the key to the present: Geological Society of London Special Publication, v. 143, p. 199–242. Weber, B., Meschede, M., Ratschbacher, L., and Frisch, W., 1997, Structure and kinematic history of the Acatlán Complex in Nuevos Horizontes–San Bernardo region, Puebla: Geofisica Internacional, v. 36, no. 2, p. 63–76. Westermann, G.E.G., Carrasco, R., and Corona, R., 1984, The Andean midJurassic Neuqueniceras ammonite assemblage of Cualac, México, in Westermann, G.E.G., ed., Jurassic–Cretaceous biochronology and biogeography of North America: St. John’s, Geological Association of Canada Special Paper 27, p. 99–112. Yañez, P., Ruiz, J., Patchett, P.J., Ortega-Gutiérrez, F., and Gehrels, G., 1991, Isotopic studies of the Acatlán Complex, southern Mexico: Implications for Paleozoic North American tectonics: Boulder, Colorado, Geological Society of America Bulletin, v. 103, p. 817–828, doi: 10.1130/00167606(1991)103<0817:ISOTAC>2.3.CO;2.
MANUSCRIPT ACCEPTED BY THE SOCIETY 3 OCTOBER 2006
Printed in the USA
Geological Society of America Special Paper 423 2007
Provenance of the Granjeno Schist, Ciudad Victoria, México: Detrital zircon U-Pb age constraints and implications for the Paleozoic paleogeography of the Rheic Ocean R. Damian Nance* Department of Geological Sciences, 316 Clippinger Laboratories, Ohio University, Athens, Ohio 45701, USA Javier Fernández-Suárez Departamento de Petrología y Geoquímica, Universidad Complutense, 28040 Madrid, Spain J. Duncan Keppie Departamento de Geología Regional, Instituto de Geología, Universidad Nacional Autónoma de México, 04510 México D.F., México Craig Storey** Department of Earth Sciences, The Open University, Walton Hall, Milton Keynes, MK7 6AA, UK Teresa E. Jeffries Department of Mineralogy, Natural History Museum, Cromwell Road, London, SW7 5BD, UK
ABSTRACT The Granjeno Schist of northeastern México is the oldest component of the Sierra Madre terrane and comprises polydeformed, pelitic metasedimentary and metavolcaniclastic rocks that enclose lenses of serpentinite-metagabbro. This low-grade Paleozoic assemblage is exposed in the core of a NNW-trending frontal anticline of the Laramide fold-thrust belt where it is tectonically juxtaposed against ca. 1 Ga granulites of the Novillo Gneiss. Silurian strata that unconformably overlie the Novillo Gneiss are unmetamorphosed and contain fauna of Gondwanan affinity. LA-ICPMS (laser ablation inductively coupled plasma mass spectroscopy)U-Pb ages for detrital zircons from a Granjeno phyllite yield age populations that cluster in the ranges ca. 1375–880 Ma, ca. 650–525 Ma, and ca. 460–435 Ma and slightly discordant grains with individual ages of ca. 1435 Ma, ca. 1640 Ma, ca. 2105 Ma, and ca. 2730 Ma. The youngest detrital zircon indicates a maximum depositional age for the Granjeno Schist of ca. 435 Ma (Lower Silurian). Detrital zircons of Neoproterozoic– Cambrian age suggest a provenance in the Maya terrane beneath the Yucatan Peninsula or the Brasiliano orogens of South America, and a source for the detrital zircons *E-mail:
[email protected]. **Present address: Department of Earth Sciences, University of Bristol, Wills Memorial Building, Queen’s Road, Bristol BS8 1RJ, UK. Nance, R.D., Fernández-Suárez, J., Keppie, J.D., Storey, C., and Jeffries, T.E., 2007, Provenance of the Granjeno Schist, Ciudad Victoria, México: Detrital zircon U-Pb age constraints and implications for the Paleozoic paleogeography of the Rheic Ocean, 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. 453–464, doi: 10.1130/2007.2423(22). For permission to copy, contact
[email protected]. ©2007 Geological Society of America. All rights reserved.
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Nance et al. of Ordovician–Silurian age is present in the Acatlán Complex of southern México. Provenance of the Mesoproterozoic detrital zircons is likely to have been the adjacent Novillo Gneiss, which has yielded ages of ca. 990–980 Ma, ca. 1035–1010 Ma, and ca. 1235–1115 Ma. These detrital ages closely match those recorded from the Cosoltepec Formation of the Acatlán Complex and support correlation of the two units, which are both interpreted to be vestiges of the southern margin of the Rheic Ocean. Keywords: Granjeno Schist, detrital zircons, Paleozoic, México, Rheic Ocean
INTRODUCTION In the southernmost Sierra Madre terrane (Fig. 1), immediately west of Ciudad Victoria in the northern Mexican state of Tamaulipas, low-grade Paleozoic metamorphic rocks of the Granjeno Schist are tectonically juxtaposed against the Novillo Gneiss, a suite of granulite-facies gneisses of Grenville (ca. 1 Ga) age. In the past, the association of these two assemblages has been linked to that of the Appalachian-Ouachita and Grenville orogens (e.g., Ortega-Gutiérrez, 1978; Ramírez-Ramírez, 1992; Ruiz et al., 1999), a correlation that would have profound implications for late Precambrian–Paleozoic continental reconstructions. However, the Novillo Gneiss is unconformably overlain by Silurian strata that contain fossils of Gondwanan rather than Laurentian affinity (Boucot et al., 1997; Stewart et al., 1999), and its correlation with the Grenville belt of Laurentia to the north is not supported by recent isotopic studies (Keppie et al., 2003; Solari et al., 2003; Cameron et al., 2004). Instead these isotopic data substantiate linkages with other outcrops of ca. 1 Ga rocks to the south, thereby supporting the hypothesis of Ortega-Gutiérrez et al. (1995) that much of eastern and southern México is underlain by a single, coherent microcontinent (Oaxaquia) of Mesoproterozoic age. The largest exposed portion of this microcontinent is the Oaxacan Complex of southern México, which is unconformably overlain by latest Cambrian–Ordovician sedimentary strata that, like the Silurian cover on the Novillo Gneiss, contain fauna of Gondwanan provenance (Robison and PantojaAlor, 1968; Ortega-Gutiérrez et al., 1995; Landing et al., 2006). Based on such linkages, the Granjeno Schist has been connected with broadly similar low-grade Paleozoic metamorphic rocks in the Acatlán Complex of the Mixteca terrane (OrtegaGutiérrez, 1978; Ramírez-Ramírez, 1992), against which the Oaxacan Complex is tectonically juxtaposed (Elías-Herrera and Ortega-Gutiérrez, 2002). Such a correlation is supported by similarities in deformational histories (Dowe et al., 2005) and would have important implications for the origin of the Granjeno Schist and the Acatlán Complex, which has been considered a vestige of the Iapetus Ocean (Ortega-Gutiérrez et al., 1999), the Rheic Ocean (Keppie and Ramos, 1999), or both (Talavera-Mendoza et al., 2005). Confirmation of such a linkage, however, is precluded by uncertainties in the age and provenance of the Granjeno Schist. In the absence of fossils, this article presents detrital zircon ages from the Granjeno Schist that provide (1) an older limit on the time of its deposition, (2) precise age constraints on
its provenance, and (3) a basis for direct comparison with rocks of the Acatlán Complex. GEOLOGICAL SETTING The Novillo Gneiss and Granjeno Schist constitute the largest exposure of Precambrian and Paleozoic rocks in northeastern México and outcrop in the core of a major basement-involved anticline (the Huizachal-Peregrina Anticlinorium) in the front ranges of the Laramide fold-thrust belt of the Sierra Madre Oriental (Fig. 2). The Novillo Gneiss comprises a variety of highgrade Mesoproterozoic rocks, whereas the Granjeno Schist is a low-grade, pre-Pennsylvanian assemblage, dominated by pelitic schist, that encloses several tectonic lenses of serpentinitemetagabbro (Ortega-Gutiérrez, 1978; Ramírez-Ramírez, 1992). Novillo Gneiss and Cover Units The Novillo Gneiss is a granulite-facies complex that includes metasedimentary rocks intruded by gabbro-anorthosite, granite, and amphibolite that have yielded ages of ca. 1035–1010 Ma and ca. 1235–1115 Ma (Cameron et al., 2004). These rocks underwent polyphase deformation and granulite-facies metamorphism at 990 ± 5 Ma, followed by post-tectonic anorthositic pegmatite emplacement at 978 ± 13 Ma (Cameron et al., 2004). Titanite and biotite from the metasedimentary rocks have yielded ages of ca. 928 Ma and ca. 697 Ma (interpreted to date cooling through ~660 °C and 300–350 °C, respectively), and an intrusive suite of unmetamorphosed northeast-trending mafic dikes has been dated at ca. 550 Ma (Keppie et al., 2006a). The Novillo Gneiss is unconformably overlain by an unmetamorphosed succession of Paleozoic marine clastics, the base of which comprises Middle Silurian (Early to Middle Wenlockian; 428–423 Ma, Gradstein et al., 2004) shallow marine strata containing fauna of Gondwanan affinity (Boucot et al., 1997; Stewart et al., 1999). This sequence is unconformably overlain by Lower Mississippian (Early Osagean; 351–342 Ma, Okulitch, 2002) sandstones and shales that contain a Laurentian shallow marine fauna suggesting deposition proximal to the North American craton. These rocks are overlain by a flow-banded rhyolite dated at 334 ± 34 Ma (lower intercept U-Pb zircon age; Stewart et al., 1999), which is unconformably overlain by Lower to Middle Pennsylvanian turbiditic lime grainstones that are, in turn, overlain by extensive Lower Permian (Wolfcampian and Leonardian;
Provenance of the Granjeno Schist, Ciudad Victoria
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Figure 1. Generalized tectonic map of México showing the distribution of terranes and their inferred boundaries (after Keppie, 2004). Oaxaquia (stippled cross-hatch) identifies that part of México underlain by ca. 1.0 Ga basement. Grey shading identifies the mainly Paleozoic rocks of the Sierra Madre and Mixteca terranes. The Granjeno Schist in the Huizachal-Peregrina Anticlinorium lies in the Sierra Madre terrane and is bound to the east by the ca. 1 Ga Novillo Gneiss (part of Oaxaquia). The Acatlán Complex forms the basement of the Mixteca terrane adjacent to Oaxaquia to the south.
Figure 2. Generalized geologic map of the Huizachal-Peregrina Anticlinorium (after Ramírez-Ramírez, 1992) showing the location of the field area.
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299–271 Ma, Gradstein et al., 2004) volcanogenic flysch (Gursky and Michalzik, 1989). All of these rocks are deformed by symmetric-asymmetric, open-close, NNW-trending folds and associated northeast-vergent thrusts (Carillo-Bravo, 1961; Gursky, 1996; Stewart et al., 1999) that predate the deposition of a Mesozoic sequence of redbeds, mafic igneous rocks, and carbonates, the oldest of which are thought to be Lower Jurassic in age (Stewart et al., 1999). This Mesozoic sequence was itself deformed during the latest Cretaceous–Middle Eocene Laramide orogeny, which produced a broad open anticline in the sub-Jurassic unconformity and ENE-vergent thin-skinned thrusting in the overlying rocks (Carillo-Bravo, 1961). Leucogranite The Novillo Gneiss is separated from the Granjeno Schist by a major NNW-trending fault along which is intruded an elongate, foliated plagioclase leucogranite (the “plagiogranite” of Ramírez-Ramírez, 1992; Fig. 2) that has been reported to be cut by ca. 546-Ma mafic dikes (Keppie et al., 2006a) and to contain xenoliths of both the Novillo Gneiss (Ortega-Gutiérrez, 1978) and the Granjeno Schist (Fries and Rincon-Orta, 1965; Gursky, 1996). This leucogranite has yielded a poorly defined U-Pb zircon lower intercept age of 351 ± 54 Ma and a 40Ar/39Ar muscovite plateau age of 313 ± 7 Ma (Dowe et al., 2005). Granjeno Schist The Granjeno Schist includes rocks of both sedimentary and igneous protoliths (Carillo-Bravo, 1961; Ortega-Gutiérrez, 1978; Ramírez-Ramírez, 1992). Metasedimentary rocks include pelitic and less common psammitic, silicic, graphitic, and volcaniclastic schist, but are typified by phyllite with bedding-parallel quartz veins rhythmically interlaminated with fine-grained psammitic schist. Mineral assemblages are dominated by quartz with phengite, chlorite (after biotite), minor albite (as porphyroblasts) and graphite, and rare pre- to syntectonic garnet (Dowe et al., 2005). Metaigneous lithologies include metabasite and tectonic lenses of serpentinite-metagabbro, the largest of which measures ~0.5 km by 10 km. The metabasite is dominated by varying amounts of actinolite, chlorite, albite, clinozoisite, epidote, calcite, and quartz. The serpentinite is a foliated lizardite chrysotile rock with occasional pyroxenite (clinopyroxene-orthopyroxene) and calcium-rich metasomatic borders. The metagabbro is massive, with a mineral assemblage dominated by actinolite, albite, chlorite, and epidote group minerals that occasionally preserve relict cumulate texture. Tectonothermal History The deformational history of the Granjeno Schist includes at least four sets of fold structures developed under greenschistfacies (350–400 °C at 300–600 MPa) metamorphic conditions (Dowe et al., 2005). The earliest deformation, D1, takes the form of rootless isoclinal structures that plunge south at moderate to
steep angles and possess a bedding-parallel axial-planar foliation. These fold sets postdate leucogranite emplacement and are considered to record phases of a single progressive deformation event under greenschist- to sub-greenschist-facies metamorphic conditions, designated D2, associated with the dextral juxapositioning of the Granjeno Schist and the Novillo Gneiss. Previous Geochronology With the exception of a 40Ar/39Ar age reported by Dowe et al. (2005), available age data for the Granjeno Schist are old and based on imprecise Rb-Sr isochrons and K/Ar thermochronology. The lower Paleozoic age traditionally assigned to the unfossiliferous Granjeno Schist is based on the oldest whole-rock Rb-Sr age of 452 ± 45 Ma (de Cserna and Ortega-Gutiérrez, 1978, recalculated for revised decay constant by Sedlock et al., 1993). However, other Rb-Sr whole-rock and whole-rock–muscovite isochron ages from the metasedimentary rocks range from 373 ± 37 Ma to 320 ± 12 Ma (de Cserna and Ortega-Gutiérrez, 1978, recalculated; Garrison et al., 1980), whereas muscovite K/Ar age data range from 318 ± 10 Ma to 257 ± 8 Ma (Fries et al., 1962; de Cserna et al., 1977), all of which suggest an episode of late Paleozoic greenschist-facies metamorphism (Ramírez-Ramírez, 1992). Dowe et al. (2005) report a 40Ar/39Ar plateau age on phengite from the Granjeno Schist of 300 ± 4 Ma (Carboniferous-Permian boundary; Gradstein et al., 2004) that they interpret to date cooling through ~350 °C (Dodson, 1973). This is similar to the 40 Ar/39Ar muscovite plateau age (313 ± 7 Ma) they obtained from the leucogranite, which they likewise interpret as dating cooling through lower greenschist temperatures. Hence they infer the tectonic juxtapositioning of the Granjeno Schist and the Novillo Gneiss, recorded by D2, to have begun in the Carboniferous and continued into the Permian. ANALYTICAL METHODS In order to clarify the age of the Granjeno Schist and better constrain its deformational history and potential regional correlation, detrital zircons were separated from a sample of Granjeno phyllite collected from the road section near the entrance to the only active serpentinite quarry (sample G-7; 23°43.15′PN, 99°16.35′W, grid reference 14QMB7222.5231.5). Zircons were separated from sample G-7 at Universidad Complutense, Madrid, following conventional techniques. Details of the separation procedure can be found in Fernández-Suárez et al. (2002) and Jeffries et al. (2003). Separated zircons were examined optically under a binocular microscope, and representative grains were picked and mounted in epoxy resin. The mount was ground down so that equatorial sections of most of the grains were exposed, and then polished to high quality. Grains were imaged by cathodoluminescence in a JEOL 5900LV scanning electron microscope at the Natural History Museum, London. Finally the mounts were cleaned by immersion in an ultrasonic bath containing a dilute HNO3 acid, and dried before being introduced to the laser ablation chamber. Analytical instrumentation
Provenance of the Granjeno Schist, Ciudad Victoria consisted of a UP213 frequency-quintupled Nd:YAG–based laser ablation system (NewWave Research, Fremont, California) coupled to a (Thermo Elemental) PQ3, quadrupole-based ICPMS (inductively coupled plasma mass spectrometer) instrument with enhanced sensitivity (S-Option) interface. The instruments, operating parameters, and procedures used for individual zircon analyses were as given in Jeffries et al. (2003). The analytical protocol and methodology, data reduction, age calculation, and common Pb correction used followed those described in Fernández-Suárez et al. (2002), Jeffries et al. (2003), and Murphy et al. (2004a,b). The samples and standard were ablated in an air-tight sample chamber flushed with helium for sample transport. The laser was focused on the sample surface, and energy density was kept constant for each analysis. In this study, the nominal laser beam diameter was 30 μm for ~65% of the analyses, but where the area to be analyzed was deemed to be large enough, a 45 μm beam was used to ensure that the analysis was performed with the optimal signal strength that the analyte volume allowed. Data were collected in discrete runs of twenty analyses comprising twelve unknowns bracketed before and after by four analyses of the standard zircon 91500 (Wiedenbeck et al. 1995). During the analysis of zircons from sample G-7, the standard 91500 yielded a weighted average of 1063.1 ± 1.6 Ma (n = 40, MSWD [mean square of weighted deviates] = 2.1) for the 206Pb/238U age (certified IDTIMS [isotopic dilution thermal ionization mass spectrometry] 206 Pb/238U age: 1062.4 ± 0.4 Ma) and a weighted average of 1064.0 ± 2.1 Ma (n = 40, MSWD = 0.7) for the 207Pb/206Pb age (certified ID-TIMS 207Pb/206Pb age: 1065.4 ± 0.3 Ma). Concordia age calculations were performed and concordia and frequency histograms/probability density distribution plots made using Isoplot v.3.00 (Ludwig, 2003). RESULTS Sixty analyses (one analysis per grain) were performed on zircons from sample G-7. Of those, four were rejected based on the presence of features such as discordance >10% (Table 1); high levels of common Pb detected in the U-Pb, Th-Pb, or PbPb isotope ratio plots; and/or elemental U-Pb fractionation or inconsistent behavior of U-Pb and Th-Pb ratios in the course of ablation (see Jeffries et al., 2003). Figure 3 shows concordia plots for the sample, with the data presented in Table 1. Where the analyses overlap concordia with a MSWD of concordance <2, we assign a U-Pb concordia age (sensu Ludwig, 2003) as the best age estimate (see bold type in Table 1). Where analyses are normally discordant (i.e., they plot below concordia), we assign the 207Pb/206Pb age and error because we are confident that any discordance is not a result of excess common Pb in the analysis or analysis-induced problems such as laser-induced elemental fractionation or mixing of differently aged domains (see Jeffries et al., 2003, for details). Consequently, these ages will approximate the “correct” age, assuming a zero-age Pb loss event, and there is a small danger that a nonzero-age thermal event could cause these
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ages to represent minimum ages. However, the amount of discordance within these zircons is minor (see Fig. 3 and Table 1), so this phenomenon is unlikely to affect any of the first-order conclusions presented herein. As shown in Figure 3 and Table 1, most analyses (~70%) of sample G-7 yielded Mesoproterozoic ages in the range of ca. 1640–1018 Ma (errors reported in Table 1), and the majority of these gave ages between ca. 1259 and 1000 Ma, with a maximum at ca. 1150 Ma (Fig. 4). Five analyses yielded earliest Neoproterozoic (Tonian) ages between ca. 991 and ca. 880 Ma, but these analyses form a continuum with the latest Mesoproterozoic zircons (Fig. 3C). Two zircons (reported in Table 1 and Fig. 4) yielded pre-Mesoproterozoic ages of ca. 2106 Ma (Rhyacian) and ca. 2736 Ma (Neoarchean). Four zircons yielded Neoproterozoic (Cryogenian–Ediacaran) ages of ca. 650, 598, and 546 Ma (two grains). Six zircons yielded Paleozoic ages ranging from ca. 527 Ma (Lower Cambrian; Gradstein et al., 2004) to ca. 433 Ma (late Llandovery; Gradstein et al., 2004). Finally, one discordant analysis corresponding to an oscillatory zoned pyramidal tip fragment yielded an imprecise 207Pb/206Pb age of 336 ± 98 Ma (Table 1). DISCUSSION Depositional Age of the Granjeno Schist Given the low-grade nature of the host pelitic rock, all of the analyzed zircons are considered detrital in origin. A maximum depositional age for the Granjeno Schist is therefore provided by the youngest precise analysis, which is only very slightly discordant and yielded a concordia age of 433 ± 14 Ma. This age is overlapped by more discordant data with concordia ages of 436 ± 6 Ma and 450 ± 7 Ma. Hence, the Granjeno Schist must be Lower Silurian or younger in age. The low precision of the youngest (discordant) analysis (336 ± 98 Ma) does not allow further refinement of the maximum depositional age, but raises the possibility that it could be significantly younger. Correlation with the Acatlán Complex A depositional age that can be no older than the Lower Silurian and could be significantly younger precludes a lower Paleozoic age for the Granjeno Schist and strengthens its correlation with similar low-grade rocks (the Cosoltepec Formation) within the Acatlán Complex of the Mixteca terrane (Fig. 1), the youngest reported detrital zircons from which are dated at ca. 455 Ma (Keppie et al., 2006b) and ca. 410 Ma (Talavera-Mendoza et al., 2005). Such a correlation has previously been proposed on the basis of lithologic similarity (Ortega-Gutiérrez, 1978; RamírezRamírez, 1992) and deformational history (Dowe et al., 2005). Both the Granjeno Schist and the Cosoltepec Formation are dominated by greenschist-facies pelitic lithologies, and both are interpreted to record two deformational episodes, the older of which is of uncertain pre-Pennsylvanian age, whereas the younger is late
Sample G-7 Anal. no. au06a06 au06a07 au06a08 au06a09 au06a11 au06a12 au06a13 au06a14 au06a15 au06a16 au06b05 au06b06 au06b07 au06b09 au06b11 au06b12 au06b13 au06b14 au06b15 au06b16 au06c05 au06c06 au06c07 au06c08 au06c09 au06c10 au06c11 au06c12 au06c13 au06c14 au06c15 au06c16 au09a05 au09a06 au09a07 au09a08 au09a09 au09a10 au09a11 au09a12 au09a13 au09a14
i.s. [s] 27 21 29 25 15 19 12 36 25 19 29 23 13 27 31 25 31 23 29 17 15 31 17 17 33 25 12 17 12 15 19 23 15 15 23 23 25 17 17 17 23 21
TABLE 1. LASER ABLATION INDUCTIVELY COUPLED PLASMA–MASS SPECTROMETER U-PB RESULTS FROM SAMPLE G-7, A PHYLLITE FROM THE GRANJENO SCHIST Ages and 2σ absolute errors Reported age* Isotopic ratios and 2σ (%) errors (Ma) 206 238 207 235 207 206 206 238 235 207 206 207 Pb/ Pb Pb/ U Pb/ U Pb/ U Pb/ U Pb/ Pb ±2σ Age (Ma) ±2σ ±2σ ±2σ ±2σ ±2σ 0.1943 1.26 2.0778 1.94 0.0776 1.64 1144 13 1142 13 1134 32 1143 0.1730 1.70 1.7961 1.54 0.0753 0.92 1028 16 1044 10 1076 18 1076 0.1460 0.68 1.3841 1.04 0.0687 1.36 879 6 882 6 890 28 880 0.1780 1.16 1.8599 1.20 0.0758 0.68 1056 11 1067 8 1088 14 1088 0.1328 1.46 1.2532 2.34 0.0684 1.80 804 11 825 13 880 38 880 0.1901 2.62 2.0258 2.00 0.0773 2.40 1122 27 1124 14 1128 48 1124 0.1748 3.16 1.8794 3.72 0.0780 1.74 1038 30 1074 25 1146 34 1146 0.1957 1.28 2.0731 2.60 0.0768 2.20 1152 13 1140 18 1116 46 1149 0.1710 1.22 1.7446 1.40 0.0740 1.82 1017 12 1025 9 1040 36 1023 0.1552 0.88 1.5108 1.32 0.0706 1.42 930 8 935 8 944 30 932 0.1634 1.40 1.6146 2.00 0.0717 2.78 976 13 976 13 976 58 976 0.2005 1.34 2.2518 1.64 0.0814 2.12 1178 14 1197 12 1232 42 1232 0.0394 1.64 0.2889 4.14 0.0532 4.40 249 4 258 9 336 98 250 0.2142 1.96 2.4859 1.90 0.0842 1.44 1251 22 1268 14 1296 28 1296 0.1979 1.22 2.1544 1.84 0.0790 1.32 1164 13 1166 13 1170 26 1165 0.1940 1.64 2.0694 2.30 0.0774 1.78 1143 17 1139 16 1130 36 1140 0.1858 0.98 1.9997 0.96 0.0781 0.82 1098 10 1115 6 1148 16 1148 0.1647 1.46 1.6617 2.24 0.0732 1.68 983 13 994 14 1018 34 1018 0.1063 1.28 0.8929 2.40 0.0609 2.48 651 8 648 12 634 54 650 0.0852 2.98 0.6820 4.48 0.0580 5.98 527 15 528 18 530 130 527 0.0821 1.64 0.6618 1.40 0.0584 1.46 509 8 516 6 546 32 546 0.1766 1.34 1.8393 2.04 0.0755 1.66 1048 13 1060 13 1082 34 1053 0.0883 0.62 0.7166 1.46 0.0589 1.24 545 3 549 6 562 26 546 0.1895 1.80 2.0985 1.44 0.0803 0.76 1118 18 1148 10 1204 16 1204 0.2300 1.34 2.6980 1.08 0.0851 1.78 1335 16 1328 8 1316 34 1327 0.1875 1.44 2.0231 2.24 0.0782 2.12 1108 15 1123 15 1152 42 1152 0.2135 2.88 2.4238 2.42 0.0823 0.92 1248 33 1250 17 1252 18 1252 0.0739 2.96 0.5872 6.48 0.0576 7.14 460 13 469 24 514 156 462 0.0695 3.68 0.5323 6.00 0.0555 8.02 433 15 433 21 432 180 433 0.2068 2.00 2.3186 1.88 0.0813 2.32 1212 22 1218 13 1228 44 1217 0.1844 1.40 2.0190 2.44 0.0794 2.38 1091 14 1122 17 1182 46 1182 0.0734 1.84 0.5645 1.80 0.0558 1.76 457 8 454 7 442 40 455 0.0720 1.72 0.5529 3.78 0.0557 3.90 448 7 447 14 440 86 450 0.0971 1.50 0.8022 2.92 0.0599 3.56 598 9 598 13 598 76 598 0.2188 0.98 2.5768 1.56 0.0854 1.26 1276 11 1294 11 1324 26 1324 0.2092 1.58 2.3830 2.14 0.0826 1.92 1224 18 1238 15 1260 38 1260 0.2079 2.62 2.4011 2.80 0.0838 1.06 1218 29 1243 20 1286 20 1286 0.2299 2.46 2.7582 2.32 0.0870 1.30 1334 30 1344 17 1360 26 1349 0.2255 1.02 2.7276 1.30 0.0877 0.66 1311 12 1336 10 1374 14 1374 0.0701 1.44 0.5247 3.92 0.0543 4.42 437 6 428 14 380 100 436 0.2901 1.64 4.0329 1.46 0.1008 1.22 1642 24 1641 12 1638 24 1641 0.1882 1.80 2.0637 2.88 0.0795 2.38 1112 18 1137 20 1184 48 1184 disc % ±2σ –0.9 12 18 4.5 1.2 5 14 2.9 38 8.6 0.5 14 34 9.4 –3.2 13 2.2 9 1.5 6 0.0 10 42 4.4 5 25.9 28 3.5 0.5 12 –1.2 15 16 4.4 34 3.4 –2.7 7 0.6 13 32 6.8 3.1 12 3.0 4 16 7.1 –1.4 8 42 3.8 0.3 11 10.5 12 –0.2 14 1.3 13 46 7.7 –3.4 6 –1.8 7 0.0 8 26 3.6 38 2.9 20 5.3 1.9 16 14 4.6 –15.0 6 –0.2 11 48 6.1 Continued
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TABLE 1. LASER ABLATION INDUCTIVELY COUPLED PLASMA–MASS SPECTROMETER U-PB RESULTS FROM SAMPLE G-7, A PHYLLITE FROM THE GRANJENO SCHIST (continued) Sample Ages and 2σ absolute errors Reported age* Isotopic ratios and 2σ (%) errors (Ma) G-7 207 207 206 206 238 235 206 238 207 235 207 206 Anal. no. i.s. [s] Pb/ Pb Pb/ U Age (Ma) disc % Pb/ U Pb/ U Pb/ U Pb/ Pb ±2σ ±2σ ±2σ ±2σ ±2σ ±2σ ±2σ au09a15 19 0.1678 1.62 1.7042 1.80 0.0737 1.44 1000 15 1010 11 1032 30 1032 30 3.1 au09a16 23 0.1659 1.00 1.6570 1.48 0.0724 0.92 989 9 992 9 996 18 0.7 991 9 au09b05 17 0.1881 1.44 2.0395 2.10 0.0786 1.36 1111 15 1129 14 1162 28 1162 28 4.4 au09b06 21 0.3814 1.62 6.8972 1.66 0.1311 0.80 2083 29 2098 15 2112 14 1.4 2106 11 au09b07 23 0.2225 1.00 2.6600 1.68 0.0867 1.94 1295 12 1317 12 1352 36 1352 36 4.2 au09b08 27 0.1825 1.84 1.9480 2.06 0.0774 0.72 1081 18 1098 14 1130 16 1130 16 4.3 au09b09 21 0.1725 2.46 1.8056 1.48 0.0759 1.70 1026 23 1047 10 1092 34 1092 34 6.0 au09b10 23 0.2476 1.54 3.1140 1.24 0.0912 1.64 1426 20 1436 10 1450 30 1.7 1435 10 au09b11 12 0.1775 1.56 1.8675 2.00 0.0763 1.36 1053 15 1070 13 1102 28 1102 28 4.4 au09b12 25 0.5247 1.06 13.6792 1.22 0.1890 0.74 2719 23 2728 12 2732 12 0.5 2731 10 au09b13 15 0.2079 0.98 2.3295 1.90 0.0813 1.56 1217 11 1221 13 1226 30 0.7 1219 11 au09b14 19 0.1887 1.04 2.0780 1.76 0.0799 2.06 1114 11 1142 12 1192 40 1192 40 6.5 au09b15 31 0.2019 1.28 2.2759 1.50 0.0818 0.94 1185 14 1205 11 1238 18 1238 18 4.3 au09b16 19 0.1676 1.40 1.7464 1.74 0.0755 0.84 999 13 1026 11 1082 16 1082 16 7.7 207 206 206 238 Notes: Abbreviations: i.s.—signal interval integrated for isotope ratio and age calculation (in seconds); disc%—percent discordance calculated from Pb/ Pb and Pb/ U ages (negative values represent reversely discordant analyses). The numbers in bold type represent concordia ages or errors sensu Ludwig (1998) only for analyses whose mean square of weighted deviates of concordance is <2. Concordia age errors include decay constant uncertainties. *See text for details.
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Paleozoic and associated with the tectonic juxtapositioning of the low-grade units against granulite-facies gneisses of ca. 1 Ga age by way of dextral ductile shear (e.g., Elías-Herrera and OrtegaGutiérrez, 2002; Malone et al., 2002; Dowe et al., 2005). In the Acatlán Complex, the younger deformational episode was a progressive event that first produced tight to isoclinal folds with a composite axial-planar foliation in response to south-vergent thrusting and dextral transpression under lower greenschistfacies conditions, and later formed upright north-south folds with an axial-planar crenulation cleavage (Malone et al., 2002). Both phases of deformation are of Early–Middle Permian age and correlate well with D2 in the Granjeno Schist, which was likewise a progressive deformational event of latest Paleozoic age and produced a bedding-subparallel composite foliation axial planar to tight to isoclinal folds followed by north-south and northeastsouthwest folds, the earlier of which possess an axial-planar crenulation cleavage (Dowe et al., 2005). The timing of the earlier (D1) deformational episode in the Granjeno Schist is constrained only by the imprecise age pre546 Ma) of the leucogranite, which is pre-tectonic with respect to D1, and by the maximum age of deposition (ca. 435 Ma). D1 in the Cosoltepec Formation was probably synchronous with Mississippian exhumation of the locally eclogitic Piaxtla Group of the Acatlán Complex, the high-grade metamorphism of which has recently been dated at 346 ± 3 Ma (Middleton et al., this volume). In the Acatlán Complex, this earlier deformational episode is considered to record exhumation (following subduction) of the high-grade Piaxtla Group onto the low-grade Cosoltepec Formation (Nance et al., this volume), a scenario that, in the Granjeno Schist, could be correlative with the initial emplacement of the serpentinite-metagabbro lenses. Further support for correlation of the Granjeno Schist and the Cosoltepec Formation is provided in this study by the age ranges of their respective detrital zircon populations (Fig. 4). Age populations in the Granjeno Schist cluster in the ranges ca. 460–435 Ma, ca. 650–525 Ma, and ca. 1375–880 Ma, whereas slightly discordant older grains show individual ages of ca. 1435 Ma, ca. 1640 Ma, ca. 2105 Ma, and ca. 2730 Ma. These closely match those recorded from the Cosoltepec Formation (Keppie et al., 2006b), which show clusters at ca. 470–455 Ma, ca. 660–500 Ma, and ca. 1340–860 Ma, and older ages of ca. 1735–1675 Ma, ca. 1880 Ma, ca. 1990 Ma, ca. 2155 Ma, and ca. 2725 Ma. This suggests that the two assemblages had similar provenances, although Mesoproterozoic zircons clearly dominate the Granjeno sample, whereas Neoproterozoic–early Paleozoic zircons dominate the Cosoltepec sample. Detrital zircon ages reported from the Cosoltepec Formation by Talavera-Mendoza et al. (2005) show broadly similar age clusters at ca. 450–410 Ma, ca. 750–550 Ma, and ca. 1000–800 Ma and minor peaks at ca. 1960 Ma, 2087 Ma, and 2197 Ma, with a similar dominance of Neoproterozoic–early Paleozoic zircons. Such a linkage, in turn, lends support to the correlation of the Novillo Gneiss and the Oaxacan Complex (e.g., Ortega-Gutiérrez et al., 1995), both of which expose rocks of similar protoliths,
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Figure 3. (A) Concordia plot of U-Pb zircon analyses of sample G-7 except one Paleoproterozoic zircon (2106 Ma) and an Archean zircon (2726 Ma), both reported in Table 1. The ellipses in all the plots represent 2σ uncertainties. (B) Enlargement of plot A showing a detail of scarce Neoproterozoic and Paleozoic zircons. (C) Detail of the dominant Mesoproterozoic zircon population.
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Figure 4. Comparative frequency and probability density distribution plots of U-Pb zircon analyses (A) from the Granjeno Schist (sample G-7; this study) and (B) from the Cosoltepec Formation (sample COS-100; Keppie et al., 2006b). All data are <10% discordant.
metamorphic grades, tectonothermal histories, and isotopic characteristics (e.g., Ruiz et al., 1999; Solari et al., 2003; Cameron et al., 2004) that are overlain by lower Paleozoic strata containing fauna of Gondwanan affinity (e.g., Robison and Pantoja-Alor, 1968; Boucot et al., 1997; Stewart et al., 1999). A prominent west to east gravity gradient from positive to negative values, which coincides with the boundary between the Paleozoic rocks and the ca. 1 Ga basement and is continuous between them (de la Fuente Duch et al., 1991), provides further support for the correlation of the lowgrade and high-grade complexes and confirms the structural evidence that the movement between them was largely along strike. Constraints on Provenance Detrital zircons from the Granjeno Schist show two major groupings of relative probability peaks: a Mesoproterozoic population that spans the interval ca. 1250–880 Ma and a late Neoproterozoic–early Paleozoic group with an age range of ca. 650–435 Ma (Fig. 3D). Potential provenances for the Mesoproterozoic detrital zircons, which constitute the dominant age population, exist within the Grenville belt of southern Laurentia to the north and in belts of similar age in Amazonia to the south (Fig. 5). However, the most likely provenance is the adjacent Novillo Gneiss, which has yielded an age range related to metamorphism and post-tectonic pegmatite emplacement of ca. 990– 980 Ma and age ranges associated with anorthosite and granitoid emplacement of ca. 1035–1010 Ma and ca. 1235–1115 Ma, respectively (Cameron et al., 2004). Almost identical age ranges of ca. 985–975 Ma, ca. 1020–1000 Ma, and ca. 1265–1100 Ma are also present in the Oaxacan Complex (Keppie et al., 2001, 2003; Solari et al., 2003), which is considered the likely provenance for the Cosoltepec Formation (Keppie et al., 2006b). That this basement source, which also includes other ca. 1 Ga inliers
in México (the Huiznopala Gneiss and Guichicovi Gneiss) with similar age ranges (Cameron et al., 2004), is of Gondwanan rather than Laurentian affinity is suggested by (1) the presence of ca. 880 Ma detrital zircons that may be found in Amazonia (Sadowski and Bettencourt, 1996, and references therein), but not Laurentia or West Africa, and (2) the paucity of detrital zircons in the Granjeno Schist and the Cosoltepec Formation that match the ca. 1.3–1.5 Ma granite-rhyolite province of Texas (Fig. 5). It is further supported by K-feldspar lead isotope data from each of these basement complexes (Cameron et al., 2004), which refute any linkage with the Laurentian basement of Texas, and by the Gondwanan fauna of the early Paleozoic cover sequences of both the Novillo Gneiss and the Oaxacan Complex. Sources for the Neoproterzoic detrital zircons may be found beneath the Yucatan Peninsula (in the Maya terrane; Fig. 1) and in the Braziliano orogens of Amazonia, but also occur to the north, in the rifted Laurentian margin of the Ouachita orogen and in the peri-Gondwanan terranes of the southern Appalachians. However, a source for the early Paleozoic zircons, while present in the Taconian belt of the southern Appalachians and the ArequipaAntofalla terrane on the southwestern margin of Amazonia, is absent in the Ouachita orogen. Given the linkages between the Granjeno Schist and the Cosoltepec Formation, however, the most likely source is the Acatlán Complex itself, the high-grade component of which (the Piaxtla Group) contains plutons (Esperanza granitoids) of ca. 480–440 Ma age (Campa-Uranga et al., 2002; Sánchez-Zavala et al., 2004; Talavera-Mendoza et al., 2005; Miller et al., this volume). Tectonic Implications Based on (1) the Silurian–Pennsylvanian age constraints for the deposition of the Granjeno Schist, (2) the apparently
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Figure 5. Detrital zircon ages from the Granjeno Schist (this article) and the Cosoltepec Formation (Keppie et al., 2006b) compared with potential source regions in Laurentia, Baltica, Amazonia, and northwest Africa (modified after Murphy et al., 2004a, and references therein). Sgp.—Subgroup; Fm.—Formation; Gps.—Groups.
distal nature of its continental-derived detritus (Dowe et al., 2005), (3) the presence within it of serpentinite-metagabbro lenses of likely oceanic origin, and (4) its potential correlation with the Cosoltepec Formation of the Acatlán Complex, we believe that the Granjeno Schist, like the Cosoltepec Formation (Keppie et al., 2006b), probably represents part of a continental rise assemblage deposited on oceanic lithosphere. If so, its depositional age range and the age populations of its detrital zircons are most simply interpreted to record its deposition on the Grenville-age Gondwanan margin of a late Paleozoic ocean. Such a margin is most readily provided by the microcontinent Oaxaquia, whose position on the southern (Gondwanan) margin of the Ordovician–Carboniferous Rheic Ocean is implied by the reconstruction (Fig. 6) of Keppie and Ramos (1999). We advocate this margin, like that inferred for the Acatlán Complex (Keppie et al., 2006b), as our preferred depositional setting of the Granjeno Schist.
CONCLUSIONS The Granjeno Schist is a polydeformed Paleozoic assemblage of low-grade, predominantly pelitic metasedimentary and metavolcaniclastic rocks that tectonically enclose lenses of serpentinite-metagabbro in the front ranges of the Laramide foldthrust belt in northeastern México. Representing the basement of the Sierra Madre terrane, these rocks are faulted against ca. 1 Ga granulite-facies gneisses (the Novillo Gneiss) that are unconformably overlain by unmetamorphosed Silurian strata containing fossils of Gondwanan affinity. Similar relationships exist in the Mixteca terrane of southern México, where Paleozoic metamorphic rocks of the Acatlán Complex, which resemble those of the Granjeno Schist in lithology and deformational history, are faulted against the Oaxacan Complex, an assemblage of ca. 1 Ga gneisses that are likewise overlain by lower Paleozoic strata of Gondwanan provenance.
Provenance of the Granjeno Schist, Ciudad Victoria Age similarities in detrital zircon populations from the Granjeno Schist and the Cosoltepec Formation of the Acatlán Complex support this linkage. Detrital zircons from the Granjeno Schist show age populations that cluster in the ranges ca. 1375–880 Ma, ca. 650– 525 Ma, and ca. 460–435 Ma, with slightly discordant grains with individual ages of ca. 1435 Ma, ca. 1640 Ma, ca. 2105 Ma, and ca. 2730 Ma. These ages, the youngest of which require that the Granjeno Schist be no older than the Lower Silurian, closely match those of the Cosoltepec Formation, which show clusters at ca. 1340–860 Ma, ca. 660–500 Ma, and ca. 470–455 Ma and older ages of ca. 1735–1675 Ma, ca. 1880 Ma, ca. 1990 Ma, ca. 2155 Ma, and ca. 2725 Ma. In both cases, the provenance of the Mesoproterozoic detrital zircons is likely to have been the adjacent ca. 1 Ga gneisses, whereas potential sources of the Neoproterozoic–Cambrian zircons lie in the Maya terrane of the Yucatan Peninsula and the Brasiliano orogens of South America. Ordovician–Silurian granitoids within the Acatlán Complex provide a likely source for the youngest detrital zircons. These data support a peri-Amazonian rather than a peri-Laurentian setting for the Granjeno Schist and are consistent with its deposition on the southern (Gondwanan) margin of the Rheic Ocean, as has previously been proposed for the Acatlán Complex. ACKNOWLEDGMENTS This project was made possible by Program for North American Mobility in Higher Education grants from the Fund for the Improvement of Postsecondary Education (to RDN) and the Secretaria de Educatión Pública de México (to JDK) and by an Ohio University 1804 Award to RDN. RDN and JDK also acknowledge the support of National Science Foundation grant EAR 0308105 and that of PAPIIT (Programa de Apoyo a Proyectos de Investigación e Innovación Tecnológica) grant IN103003 and CONACyT (Consejo Nacional de Ciencia y Tecnología) funding, respectively. Thoughtful reviews by Quentin Crowley and JeanPaul Liégeois greatly improved the article, which is a contribution to IGCP (International Geoscience Program) Project 497: The Rheic Ocean: Its Origin, Evolution, and Correlatives. REFERENCES CITED Boucot, A.J., Blodgett, R.B., and Stewart, J.H., 1997, European Province Late Silurian brachiopods from the Ciudad Victoria area, Tamaulipas, northeastern Mexico, in Klapper, G., Murphy, M.A., and Talent, J.A., eds., Paleozoic sequence stratigraphy, biostratigraphy, and biogeography: Studies in honor of J. Granville (“Jess”) Johnson: Boulder, Colorado, Geological Society of America Special Paper 321, p. 273–293. Cameron, K.L., Lopez, R., Ortega-Gutiérrez, F., Solari, L.A., Keppie, J.D., and Schulze, C., 2004, U-Pb constraints and Pb isotopic compositions of leached feldspars: Constraints on the origin and evolution of Grenvillian rocks from eastern and southern Mexico, in Tollo, R.P., et al., eds., Proterozoic tectonic evolution of the Grenville orogen in North America: Boulder, Colorado, Geological Society of America Memoir 197, p. 755–770. Campa-Uranga, M.F., Gehrels, G., and Torres de León, R., 2002, Nuevas edades de granitoides metamorfizados del Complejo Acatlán en el Estado de Guerrero: Actas Instituto Nacional de Geoquímica, v. 8, no. 1, p. 248.
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Avalonia 550-750 Ma
Cadomia BALTICA
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Yucatan basement ca. 540-560 Ma Amazon craton AFRICA
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SOUTH AMERICA Arequipa Cuyania
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(ca. 420 Ma) Figure 6. Silurian (ca. 420 Ma) continental reconstruction modified from Keppie (2004) showing the locations of Oaxaquia and the Granjeno and Cosoltepec continental rises on the Gondwana margin of the Rheic Ocean.
Carillo-Bravo, J., 1961, Geología del anticlinoria Huizachal-Peregrina al NW de Ciudad Victoria, Tamaulipas: Asociación Mexicana de Geólogos Petroleros Boletín, v. 13, p. 1–98. de Cserna, Z., and Ortega-Gutiérrez, F., 1978, Reinterpretación tectónica del Equisto Granjeno de ciudad Victoria, Tamaulipas: Contestación: Universidad Nacional Autónoma de México, Instituto de Geología, Revista, v. 2, p. 212–215. de Cserna, Z., Graf, J.L., Jr., and Ortega-Gutiérrez, F., 1977, Aloctono del Paleozoico inferior en la region de Ciudad Victoria: Estado de Tamaulipas: Universidad Nacional Autónoma de México, Instituto de Geología, Revista, v. 1, p. 33–43. de la Fuente Duch, M.F., Aiken, C.L.V., and Manuel Mena, J., 1991, Cartas gravimetricas de la Republica Mexicana: Universidad Nacional Autónoma de México, Instituto de Geofisica, scale 1:3,000,000. Dodson, M.H., 1973, Closure temperature in cooling geochronological and petrological systems: Contributions to Mineralogy and Petrology, v. 40, p. 259–274, doi: 10.1007/BF00373790. Dowe, D.S., Nance, R.D., Keppie, J.D., Cameron, K.L., Ortega-Rivera, A., Ortega-Gutiérrez, F., and Lee, J.W.K., 2005, Deformational history of the Granjeno Schist, Ciudad Victoria, Mexico: Constraints on the closure of the Rheic Ocean?: International Geology Review, v. 47, p. 920–937. Elías-Herrera, M., and Ortega-Gutiérrez, F., 2002, Caltepec fault zone: An Early Permian dextral transpressional boundary between the Proterozoic Oaxacan and Paleozoic Acatlán complexes, southern Mexico, and regional implications: Tectonics, v. 21, p. 1–18, doi: 10.1029/2000TC001278. Fernández-Suárez, J., Gutiérrez-Alonso, G., and Jeffries, T.E., 2002, The importance of along-margin terrane transport in northern Gondwana: Insights from detrital zircon parentage in Neoproterozoic rocks from Iberia and Brittany: Earth and Planetary Science Letters, v. 204, p. 75–88, doi: 10.1016/S0012-821X(02)00963-9. Fries, C., and Rincon-Orta, C., 1965, Nuevas aportaciones geocronológicas y tecnicas empleadas en al laboratorio de geocronología, Universidad Nacional Autónoma de México, Instituto de Geología: Boletin (Instituto de Estudios de Poblacion y Desarrollo [Dominican Republic]), v. 73, p. 57–134. Fries, C., Jr., Schmitter, E., Damon, P.E., Livingston, D.E., and Erikson, R., 1962, Edad de las rocas metamorficas en los Canones de la Peregrina y de Caballeros, parte centro-occidental de Tamaulipas, Universidad Nacional
464
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Autónoma de México, Instituto de Geologia: Boletin (Instituto de Estudios de Poblacion y Desarrollo [Dominican Republic]), v. 64, no. 4, p. 55–59. Garrison, J.R., Ramírez-Ramírez, C., and Long, L.E., 1980, Rb/Sr isotopic study of the ages and provenance of Precambrian granulite and Paleozoic greenschist near Ciudad Victoria, Mexico, in The origin of the Gulf of Mexico and the early opening of the Central North Atlantic Ocean: Baton Rouge, Louisiana State University, p. 37–49. Gradstein, F.M., Ogg, J.G., Smith, A.G., Bleeker, W., and Lourens, L.J., 2004, A new geologic time scale, with special reference to Precambrian and Neogene: Episodes, v. 27, p. 83–100. Gursky, H.-J., 1996, Paleozoic stratigraphy of the Peregrina Canyon area, Sierra Madre Oriental: Zentralblatt für Geologie und Paläontologie: Teil I, v. 7/8, p. 973–989. Gursky, H.-J., and Michalzik, D., 1989, Lower Permian turbidites in the northern Sierra Madre Oriental, México: Zentralblatt für Geologie und Paläontologie: Teil I, v. 5/6, p. 821–838. Jeffries, T., Fernández-Suárez, J., Corfu, F., and Gutiérrez Alonso, G., 2003, Advances in U-Pb geochronology using a frequency quintupled Nd:YAG based laser ablation system (lambda = 213 nm) and quadrupole based ICP-MS: Journal of Analytical Atomic Spectrometry, v. 18, p. 847–855, doi: 10.1039/b300929g. Keppie, J.D., 2004, Mexican terranes revisited: A 1.3 Ga odyssey: International Geology Review, v. 46, p. 765–794. Keppie, J.D., and Ramos, V.S., 1999, Odyssey of terranes in the Iapetus and Rheic Oceans during the Paleozoic, in Keppie, J.D., and Ramos, V.A., eds., Laurentia-Gondwana connections before Pangea: Boulder, Colorado, Geological Society of America Special Paper 336, p. 267–276. Keppie, J.D., Dostal, J., Ortega-Gutiérrez, F., and Lopez, R., 2001, A Grenvillian arc on the margin of Amazonia: Evidence from the southern Oaxacan Complex, southern Mexico: Precambrian Research, v. 112, p. 165–181, doi: 10.1016/S0301-9268(00)00150-9. Keppie, J.D., Dostal, J., Cameron, K.L., Solari, L.A., Ortega-Gutiérrez, F., and Lopez, R., 2003, Geochronology and geochemistry of Grenvillian igneous suites in the northern Oaxacan Complex, southern Mexico: Tectonic implications: Precambrian Research, v. 120, p. 365–389, doi: 10.1016/ S0301-9268(02)00166-3. Keppie, J.D., Dostal, J., Nance, R.D., Miller, B.V., Ortega-Rivera, A., and Lee, J.K.W., 2006a, Circa 546 Ma plume-related dykes in the ~1 Ga Novillo Gneiss (east-central Mexico): Evidence for the initial separation of Avalonia: Precambrian Research, v. 147, p. 342–353. Keppie, J.D., Nance, R.D., Fernández-Suárez, J., Jeffries, T.E., Storey, C., and Murphy, J.B., 2006b, Detrital zircon data from the eastern Mixteco terrane: Evidence for an Ordovician-Mississippian continental rise and a Permo-Triassic clastic wedge adjacent to Oaxaquia: International Geology Review, v. 48, p. 97–111. Landing, E., Keppie, J.D., and Westrop, S.R., 2006, Lower Paleozoic of northwest Gondwana: Terminal Cambrian and lowest Ordovician Tiñu Formation in southern Mexico: Geological Society of America, Abstracts with Programs, v. 38, no. 2, p. 21. Ludwig, K.R., 1998, On the treatment of concordant uranium-lead ages: Geochimica et Cosmochimica Acta, v. 62, p. 665–676, doi: 10.1016/S00167037(98)00059-3. Ludwig, K.R., 2003, Isoplot 3.00: Berkeley Geochronology Center Special Publication 4, 70 p. Malone, J.W., Nance, R.D., Keppie, J.D., and Dostal, J., 2002, Deformational history of part of the Acatlán Complex: Late Ordovician–Early Silurian and Early Permian orogenesis in southern Mexico: Journal of South American Earth Sciences, v. 15, p. 511–524, doi: 10.1016/S08959811(02)00080-9. Middleton, M., Keppie, J.D., Murphy, J.B., Miller, B.V., Nance, R.D., OrtegaRivera, A., and Lee, J.K.W., 2007 (this volume), P-T-t constraints on exhumation following subduction in the Rheic Ocean from eclogitic rocks in the Acatlán Complex of southern México, in Linnemann, U., et al., eds., The evolution of the Rheic Ocean: From Avalonian-Cadomian active margin to Alleghenian-Variscan collision: Boulder, Colorado, Geological Society of America Special Paper 423, doi: 10.1130/2007.2423(25). Miller, B.V., Dostal, J., Keppie, J.D., Nance, R.D., Ortega-Rivera, A., and Lee, J.W.K., 2007 (this volume), 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, in Linnemann, U., et al., eds., The evolution of the Rheic Ocean: From Avalonian-Cadomian active margin to Alleghenian-Variscan collision: Boulder, Colorado, Geological Society of America Special Paper 423, doi: 10.1130/2007.2423(23).
Murphy, J.B., Fernández-Suárez, J., and Jeffries, T., 2004a, Lithogeochemical, Sm-Nd and U-Pb isotopic data from the Silurian–Early Devonian Arisaig Group clastic rocks, Avalon terrane, Nova Scotia: A record of terrane accretion in the Appalachian-Caledonide orogen: Geological Society of America Bulletin, v. 116, p. 1183–1201, doi: 10.1130/B25423.1. Murphy, J.B., Fernández-Suárez, J., Jeffries, T., and Strachan, R.A., 2004b, U-Pb (LA-ICP-MS) dating of detrital zircons from Cambrian clastic rocks of Avalonia: Erosion of a Neoproterozoic arc along the northern Gondwanan margin: Journal of the Geological Society of London, v. 161, p. 243–254. Nance, R.D., Miller, B.V., Keppie, J.D., Murphy, J.B., and Dostal, J., 2007 (this volume), Vestige of the Rheic Ocean in North America: The Acatlán Complex of southern México, in Linnemann, U., et al., eds., The evolution of the Rheic Ocean: From Avalonian-Cadomian active margin to Alleghenian-Variscan collision: Geological Society of America Special Paper 423, doi: 10.1130/2007.2423(21). Okulitch, A.V., 2002, Geological time scale 2003: Ottawa, Ontario, Geological Survey of Canada, Open File 3040, National Earth Science Series, Geological Atlas—Revision. Ortega-Gutiérrez, F., 1978, El Gneiss Novillo y rocas metamorficas asociadas en los canones del Novillo y la Peregrina, area Ciudad Victoria, Tamaulipas: Universidad Nacional Autónoma de México, Instituto de Geología, Revista, v. 2, p. 19–30. Ortega-Gutiérrez, F., Ruiz, J., and Centeno-Garcia, E., 1995, Oaxaquia, a Proterozoic microcontinent accreted to North America during the late Paleozoic: Geology, v. 23, p. 1127–1130, doi: 10.1130/0091-7613(1995)023<1127: OAPMAT>2.3.CO;2. Ortega-Gutiérrez, F., Elías-Herrera, M., Reyes-Salas, M., Macias-Romo, C., and Lopez, R., 1999, Late Ordovician–Early Silurian continental collisional orogeny in southern Mexico and its bearing on GondwanaLaurentia connections: Geology, v. 27, p. 719–722, doi: 10.1130/00917613(1999)027<0719:LOESCC>2.3.CO;2. Ramírez-Ramírez, C., 1992, Pre-Mesozoic geology of Huizachal-Peregrina Anticlinorium, Ciudad Victoria, Tamaulipas, and adjacent parts of eastern Mexico: [Ph.D. thesis]: Austin, University of Texas, 317 p. Robison, R., and Pantoja-Alor, J., 1968, Tremadocian trilobites from Nochixtlan region, Oaxaca, Mexico: Journal of Paleontology, v. 42, p. 767–800. Ruiz, J., Tosdal, R.M., Restrepo, P.A., and Murillo-Muñetón, G., 1999, Pb isotope evidence for Colombia–southern México connections in the Proterozoic, in Keppie, J.D., and Ramos, V.A., eds., Laurentia-Gondwana connections before Pangea: Boulder, Colorado, Geological Society of America Special Paper 336, p. 183–197. Sadowski, G.R., and Bettencourt, J.S., 1996, Mesoproterozoic tectonic correlations between eastern Laurentia and the western border of the Amazon craton: Precambrian Research, v. 76, p. 213–227, doi: 10.1016/0301– 9268(95)00026–7. Sánchez-Zavala, J.L., Ortega-Gutiérrez, F., Keppie, J.D., Jenner, G.A., Belousova, E., and Maciás-Romo, C., 2004, Ordovician and Mesoproterozoic zircons from the Tecomate Formation and Esperanza Granitoids, Acatlán Complex, southern Mexico: Local provenance in the Acatlán and Oaxacan complexes: International Geology Review, v. 46, p. 1005–1021. Sedlock, R.L., Ortega-Gutiérrez, F., and Speed, R.C., 1993. Tectonostratigraphic terranes and tectonic evolution of Mexico: Boulder, Colorado, Geological Society of America Special Paper 278, 146 p. Solari, L.A., Keppie, J.D., Ortega-Gutiérrez, F., Cameron, K.L., Lopez, R., and Hames, W.E., 2003, 990 and 1100 Ma Grenvillian tectonothermal events in the northern Oaxacan Complex, southern Mexico: Roots of an orogen: Tectonophysics, v. 365, p. 257–282, doi: 10.1016/S0040-1951(03)00025-8. Stewart, J.H., Blodgett, R.B., Boucot, A.J., Carter, J.L., and Lopez, R., 1999, Exotic Paleozoic strata of Gondwanan provenance near Ciudad Victoria, Tamaulipas, Mexico, in Ramos, V.A., and Keppie, J.D., eds., LaurentiaGondwana connections before Pangea: Boulder, Colorado, Geological Society of America Special Paper 336, p. 227–252. Talavera-Mendoza, O., Ruiz, J., Gehrels, G.E., Meza-Figueroa, D.M., VegaGranillo, R., and Campa-Uranga, M.F., 2005, U-Pb geochronology of the Acatlán Complex and implications for the Paleozoic paleogeography and tectonic evolution of southern Mexico: Earth and Planetary Science Letters, v. 235, p. 682–699, doi: 10.1016/j.epsl.2005.04.013. Wiedenbeck, M., Allé, P., Corfu, F., Griffin, W.L., Meier, M., Orbeli, F., von Quadt, A., Roddick, J.C., and Spiegel, W., 1995, Three natural zircon standards for U-Th-Pb, Lu-Hf, trace element and REE analyses: Geostandards Newsletter, v. 19, p. 1–23. MANUSCRIPT ACCEPTED BY THE SOCIETY 3 OCTOBER 2006 Printed in the USA
Geological Society of America Special Paper 423 2007
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 Brent V. Miller* Department of Geology and Geophysics, Texas A&M University, College Station, Texas 77843, USA Jaroslav Dostal Department of Geology, St. Mary’s University, Halifax, Nova Scotia, Canada B3H 3C3 J. Duncan Keppie Departamento de Geología Regional, Instituto de Geología, Universidad Nacional Autónoma de México, 04510 Mexico D.F. R. Damian Nance Department of Geological Sciences, Ohio University, Athens, Ohio, 45701 USA Amabel Ortega-Rivera Estación Regional del Noroeste, Instituto de Geología, Universidad Nacional Autónoma de México, Hermosillo, Sonora, México James K.W. Lee Department of Geological Sciences and Geological Engineering, Queens University, Kingston, Ontario, Canada K7L 3N6
ABSTRACT U-Pb zircon data from three undeformed to slightly deformed, megacrystic, granitoid plutons in the northern Acatlán Complex of southern México has indicated that all three are part of a larger suite of late Ordovician plutons. 40Ar/39Ar data from hornblende and biotite show mainly disturbed spectra, but biotite from the Palo Liso and Los Hornos plutons yields plateaus with ages of 305 ± 26 Ma and 157 ± 12 Ma, respectively. These thermal events may be correlated, respectively, with Permo-Triassic and Jurassic tectonothermal events recorded elsewhere in the Acatlán Complex. All three plutons are peraluminous with calc-alkaline affinities, characteristics that are consistent with inherited zircon ages and together suggest a source in Mesoproterozoic calc-alkaline rocks similar to those exposed in the neighboring Oaxaca terrane. We interpret these granites to be related to the early Ordovician separation of *E-mail:
[email protected]. Miller, B.V., Dostal, J., Keppie, J.D., Nance, R.D., Ortega-Rivera, A., and Lee, J.K.W., 2007, 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, 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. 465–475, doi: 10.1130/2007.2423(23). For permission to copy, contact
[email protected]. ©2007 Geological Society of America. All rights reserved.
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Miller et al. peri-Gondwanan terranes from Gondwana during the opening of the Rheic Ocean. Elsewhere in the Acatlán Complex, Ordovician megacrystic granitoids of the Piaxtla Suite were subjected to high-grade metamorphism, which we infer to be related to subduction along the Gondwanan margin during the Devonian–Carboniferous. The three plutons reported here were not affected by Devono-Carboniferous metamorphism and thus are inferred to have remained outside the subduction zone. Keywords: Acatlán Complex, geochronology, geochemistry, Rheic Ocean, Gondwana
INTRODUCTION The Acatlán Complex of southern México (Fig. 1; OrtegaGutiérrez, 1978) represents the largest exposure of Paleozoic rocks in the country and forms the basement to the Mixteca terrane (Campa and Coney, 1983). The complex comprises an assemblage of multiply deformed metasedimentary units, locally eclogitic mafic-ultramafic suites, and K-feldspar megacrystic granitoid bodies. The Acatlán Complex is tectonically juxtaposed to the east against ca. 1 Ga granulite-facies gneisses of the Oaxaca Complex (Fig. 1; Keppie et al., 2003a; Solari et al., 2003), the early Paleozoic cover of which contains Gondwanan fauna (Robison and Pantoja-Alor, 1968). The inferred tectonostratigraphic setting and previous geochronology suggested a punctuated early Paleozoic history for much of the Acatlán Complex, with Ordovician–Silurian (Acatecan), Devonian (Mixtecan), and Permian deformation and plutonic events—a history similar to the Taconian-Acadian-Alleghanian, Iapetus-related history of the southern Appalachian orogen (e.g., Ortega-Gutiérrez et al., 1999). Recent evidence, however, points to an origin in the Rheic Ocean (Nance et al., this volume, Chapter 21). Ordovician to early Silurian variably deformed, commonly megacrystic, calcalkaline granitoids with mixed arc and withinplate affinities form part of a bimodal suite in the Acatlán Complex. The megacrystic granitoids have played an important role both in the previous tectonostratigraphic interpretation of the Acatlán Complex and in its recent reassessment because they were inferred to result from the main orogenic event that shaped the complex (Ortega-Gutiérrez et al., 1999). Our mapping, dating, and geochemical and structural analysis of the Acatlán Complex led us to question previous interpretations of the age and field relations of some granitoids (Nance et al., this volume, Chapter 21). As part of this reassessment we present here new geochemical and geochronologic data from three undeformed to slightly deformed Ordovician megacrystic plutons in the northern part of the Acatlán Complex—the La Noria, Los Hornos, and Palo Liso plutons. GEOLOGIC SETTING AND PREVIOUS GEOCHRONOLOGY Megacrystic granitoids in the Acatlán Complex were previously combined into a single map unit termed “Esperanza granitoids” (Ortega-Gutiérrez, 1981), which yielded Ordovician RbSr ages (Fries and Rincon-Orta, 1965; Halpern et al., 1974). The
presence of high-silica phengite, grossular-rich garnet, and other petrographic features (Ortega-Gutiérrez et al., 1999) was thought to provide a link, through shared high-pressure metamorphism, between the granitoids and eclogites in a unit now termed the Piaxtla Group. Both the granitoids and the eclogite-bearing unit were involved in the main deformational and metamorphic event that assembled the Acatlán Complex, which was at that time considered Late Ordovician to Early Silurian based on one relatively undeformed granitoid exposure near La Noria (Figs. 1 and 2A) with a U-Pb zircon lower intercept age of 371 ± 34 Ma (Yañez et al., 1991). Another highly deformed megacrystic granitoid from the Esperanza type locality (Figs. 1 and 2B) yielded a lower intercept age of 440 ± 14 Ma (Ortega-Gutiérrez et al., 1999). This led to subdivision of the Esperanza granitoids into (1) strongly deformed Ordovician–Silurian Esperanza-type granitoids that are inferred to have resulted from decompression melting following eclogitefacies metamorphism during the Acatecan orogeny (Ortega-Gutiérrez et al., 1999; Ramirez-Espinosa, 2001) and (2) undeformed or slightly deformed Devono-Carboniferous granitoids developed as part of a magmatic arc during the mainly greenschist-facies Mixtecan orogeny (Sánchez-Zavala et al., 2000). Recent dating by Sánchez-Zavala et al. (2004) and TalaveraMendoza et al. (2005) has resulted in eight new Ordovician to early Silurian ages from Esperanza granitoids throughout the Acatlán Complex. Two other granitoids analyzed by TalaveraMendoza et al. (2005) were interpreted to yield Mesoproterozoic crystallization ages, but the interpretation of these data is not unique, and additional work will be required to substantiate the claim of Mesoproterozoic plutonism in the Acatlán Complex. Other recent work has shown that the eclogite-facies metamorphism and decompression migmatites that were thought to be the melt source for the Esperanza granitoids (Ortega-Gutiérrez et al., 1999) are not Ordovician, but early Carboniferous (Middleton et al., this volume), and that the Mixtecan orogenic event is PermoTriassic (Malone et al., 2002; Keppie et al., 2004a, 2006). Such changes require reassessment of the age, geochemistry, and tectonic setting of the megacrystic granitoids. Megacrystic granitoids are characterized by a 1–6 cm pink K-feldspar crystals (Fig. 2) set in a coarse matrix of quartz, plagioclase, K-feldspar, and biotite with rare hornblende and accessory zircon and Fe-Ti oxide minerals. Chlorite and sericite are common secondary minerals. More detailed field relations and petrography can be found in Yañez et al. (1991), Ortega-Gutiérrez et al. (1999), and Ramirez-Espinosa (2001).
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Xayacatlán Fm. Ordovician Granitoid plutons Precambrian ca. 1 Ga Oaxacan Complex
Acatlán Complex
Post-Permian cover Permo-Triassic Chazumba Petlalcingo Lithodeme Magdalena Suite Lithodeme Mississippian-Middle Permian
Figure 1. Geological map of the Acatlán Complex, southern México, showing sample locations (modified after Ortega-Gutiérrez et al., 1999). Fm.—Formation.
468
Miller et al. NEW U-Pb GEOCHRONOLOGY Analytical Techniques We sampled three megacrystic granitoids in the northern part of the Acatlán Complex (Fig. 1) to attempt to better define their crystallization and postmetamorphic cooling ages and to more fully describe their geochemistry. The Los Hornos granitoid (LH-1) pluton contains K-feldspar megacrysts up to 4 cm in diameter, is not penetratively foliated, and is unconformably overlain by conglomerate and limestone of the Middle Permian Tecomate Formation (Keppie et al., 2004a). Immediately to the south is the La Noria granitoid pluton (LN-1; Figs. 1 and 2A), which is foliated and cuts psammitic metasedimentary rocks previously correlated with the Tecomate Formation. Given the recent reevaluation of the age of both the Tecomate Formation and the La Noria granite (Nance et al., this volume, Chapter 21), the stratigraphic relationships of these psammitic metasediments to other units in the complex are currently unclear. The La Noria and Los Hornos bodies are possibly continuous beneath post-Permian cover sediments. The Palo Liso pluton (PL-1; Figs. 1 and 2C) is mildly deformed and is faultbounded against rocks of the uppermost Devonian Otates Formation (Vachard et al., 2000; Derycke-Khatir et al., 2005). Lithologically similar megacrystic granitoids also occur in the Piaxtla Suite; however, they are generally intensely deformed and have been subjected to high-grade metamorphism (Reyes-Salas, 2003). The samples were crushed and zircons were separated, processed, and analyzed by isotope dilution thermal ionization mass spectrometry (ID-TIMS) in the Radiogenic Isotope Geochemistry Laboratory at the University of North Carolina–Chapel Hill using the techniques described in Miller et al. (2006). Zircon grains were carefully hand-picked in an effort to select clear, wellfaceted, inclusion- and core-free grains. Most zircon analyses were conducted on single grains or on a single fragment broken from a grain; the smallest zircons required fractions of 2–5 grains (Table 1). In addition, selected zircon grains from each sample were analyzed by SHRIMP-RG (sensitive high-resolution ion microprobe–reverse geometry) at Stanford University following techniques described in DeGraaff-Surpless et al. (2002). Uranium and thorium concentrations were measured relative to the CZ3 zircon standard, and isotopic ratios and concentrations were measured relative to the R33 zircon standard from a quartz diorite of the Braintree Complex, Vermont. Data were reduced with SQUID version 1.02 program of Ludwig (2001) and plotted with Isoplot Ex, version 2.49 (Ludwig, 2003). See notes to Table 1 for additional analytical details. Results
Figure 2. Field examples of megacrystic granitoids: (A) granitoid from the La Noria pluton, (B) deformed “Esperanza” granitoid from the type locality northwest of Acatlán, and (C) granitoid from the Palo Liso pluton. For scale, the Mexican $10 coin is 28 mm in diameter.
Palo Liso granite zircons contain common Mesoproterozoic inheritance, recognized in the ID-TIMS data, despite many single-grain analyses, as a highly discordant trend with an upper intercept of ca. 1065 Ma and a lower intercept similar to the granite crystallization age (Fig. 3A) and in the SHRIMP data by
TABLE 1. U-Pb ISOTOPIC DATA FOR ACATLÁN COMPLEX ORDOVICIAN GRANITES Total§ Total§ Atomic ratios Total† 206 206 207 207 206 Analysis no., fraction Wt. U Pb Common Pb U Pb Pb# % % Pb§ Pb# Pb# Pb# % 204 208 238 206 235 (number of grains) Pb U Error 4 U Error†† (ng) (pg) (pg) (ppm) (ppm) Pb Pb Error†† (mg)† Palo Liso granite: ID-TIMS (PL-1; 18°30.667′, 98°17.778′) 1, long thin prism (1) 0.001 1.50 104.1 5.78 1505 104 1168 15.698 0.06989 0.439 0.54189 0.457 0.05624 0.127 2, medium prism (1) 0.001 0.90 59.6 1.49 896 60 2705 21.999 0.07029 0.123 0.54513 0.150 0.05625 0.084 3, long thin prism (1) 0.001 2.01 141.0 1.51 2009 141 6001 10.056 0.07060 0.062 0.54710 0.122 0.05621 0.104 4, medium prism (1) 0.001 2.86 202.9 12.37 2859 203 1061 24.795 0.07171 0.255 0.56991 0.281 0.05764 0.113 5, medium prism (1) 0.001 1.04 76.73 6.49 1037 77 752 13.613 0.07270 0.634 0.56341 0.646 0.05621 0.123 6, large stubby prism (1) 0.003 0.98 76.69 2.24 328 26 2104 6.225 0.07444 0.113 0.57718 0.169 0.05623 0.120 7, small prisms (2) 0.002 1.23 90.59 1.53 613 45 3910 14.012 0.07620 0.082 0.59788 0.361 0.05691 0.333 8, medium prism (1) 0.001 1.31 104.0 4.79 1305 104 1432 18.731 0.08185 0.437 0.67179 0.449 0.05953 0.099 9, thin prism (1) 0.001 1.53 142.7 1.99 1528 143 4719 15.251 0.09626 0.078 0.82838 0.101 0.06241 0.063 10, large stubby prism (1) 0.002 1.64 152.2 1.89 818 76 5329 16.634 0.09651 0.091 0.82550 0.186 0.06204 0.155 11, small equant (2) 0.002 1.92 184.5 0.90 959 92 13481 15.285 0.09897 0.101 0.87822 0.136 0.06436 0.091 12, medium prism (1) 0.001 0.44 47.64 1.82 436 48 1745 17.093 0.11303 0.131 1.01771 0.180 0.06530 0.120 13, medium prism (1) 0.001 1.29 165.5 1.16 1293 166 9156 11.515 0.12860 0.268 1.25005 0.312 0.07050 0.160 14, large stubby prism (1) 0.003 7.28 1019 1.76 2427 340 35680 7.836 0.13512 0.080 1.42985 0.097 0.07675 0.054 La Noria granite: ID-TIMS (LN-1; 18°24′, 98°11′30′′) 1, medium prism (1) 0.001 1.44 137.3 3.83 1441 137 2380 16.676 0.09858 0.200 0.87543 0.216 0.06441 0.080 2, thin prism (1) 0.001 1.84 194.8 4.76 1840 195 2654 14.358 0.10793 0.349 0.99293 0.383 0.06673 0.153 3, medium prism (1) 0.001 1.84 234.2 5.27 1840 234 2824 11.739 0.12756 0.280 1.20743 0.291 0.06865 0.080 4, large stubby prism (1) 0.001 4.67 729.5 5.33 4674 730 8669 13.192 0.15661 0.127 1.74708 0.140 0.08091 0.059 Atomic ratios 206 206 207 207 U Th Pb* Pb# % Pb# % Pb# % 238 206 235 Sample-spot no. (ppm) (ppm) (ppm) U Error## U Error 6 Pb Error## Palo Liso granite: SHRIMP (PL-1; 18°30.667′, 98°17.778′) PL1-1 813 77.40 36.7 0.05250 2.7 0.50900 9.4 0.0703 9.1 PL1-3 363 33.24 23.0 0.07345 0.5 0.57900 2.0 0.0572 1.9 PL1-4 159 83.21 10.7 0.07800 3.3 0.60000 3.9 0.0558 2.2 PL1-5 79 22.53 14.0 0.20660 3.3 2.28300 3.7 0.0802 1.6 PL1-6 608 65.09 42.2 0.08060 3.2 0.61500 3.9 0.0553 2.1 PL1-7 272 109.09 17.8 0.07570 3.2 0.56100 4.9 0.0538 3.6 PL1-8 74 31.06 4.9 0.07590 3.4 0.56600 5.8 0.0541 4.8 La Noria granite: SHRIMP (LN-1; 18°24′, 98°11′30′′) LN1-1 174 60.44 30.2 0.20240 1.0 2.26500 1.4 0.08117 10.0 LN1-2 244 76.09 42.4 0.20240 0.9 2.20500 1.3 0.07902 0.9 LN1-3 203 45.22 13.2 0.07613 1.1 0.63200 2.9 0.0602 2.6 LN1-4 22 7.99 3.6 0.19200 5.5 1.72000 11.0 0.0647 9.9 LN1-5 220 56.68 37.3 0.19730 1.0 2.14700 1.6 0.0789 1.2 LN1-6 402 305.21 25.8 0.07444 1.0 0.57600 2.3 0.0562 2.1 LN1-7 194 12.26 12.5 0.07478 1.1 0.54500 3.8 0.0528 3.7 LN1-8 984 539.41 64.3 0.07613 0.9 0.60760 1.3 0.05788 1.0 LN1-9 293 25.86 19.3 0.07631 1.0 0.55600 2.9 0.0528 2.7 LN1-10 232 31.29 15.5 0.07740 1.4 0.57800 3.3 0.0542 3.0 LN1-11 208 67.09 36.1 0.20260 1.0 2.24800 1.5 0.08049 1.2 LN1-12 269 129.20 60.4 0.26150 1.1 3.42400 1.3 0.09496 0.8 LN1-13 641 51.82 40.2 0.07303 0.9 0.56510 1.6 0.05612 1.3 LN1-14 1217 127.01 171.0 0.16300 0.9 1.69400 1.1 0.07538 0.6 Los Hornos Granite: SHRIMP (LH-1; 18°24′, 98°11′30′′) LH1-1 311 82.74 52.6 0.19670 1.2 2.16600 1.5 0.07986 0.9 LH1-2 432 43.79 25.6 0.06857 0.6 0.50200 2.1 0.0531 2.0 LH1-3 537 69.40 32.8 0.07105 0.7 0.54530 1.6 0.05566 1.4 3 † Weight estimated from measured grain dimensions and assuming that density = 4.67g/cm , ~20% uncertainty affects only U and Pb concentrations. § Corrected for fractionation (0.12 ± 0.08%/amu, Faraday-Daly; 0.20 ± 0.1%/amu, Daly) and spike. # Corrected for fractionation, blank, and initial common Pb. †† ID-TIMS (isotope dilution thermal ionization mass spectrometry) errors are 2σ. §§ 207 Pb/235U-206Pb/238U correlation coefficient of Ludwig (1989). ## SHRIMP (sensitive high-resolution ion microprobe) errors are 1σ. Ages (Ma) 207 Pb Pb 235 206 U Pb
439.7 461.7 441.8 462.2 443.1 460.6 458.0 516.2 453.7 460.5 462.6 461.6 475.9 487.9 521.8 586.5 612.7 688.3 611.1 675.4 640.0 753.4 712.8 784.0 823.4 942.9 901.5 1114.7
435.5 437.9 439.7 446.4 452.4 462.9 473.4 507.1 592.5 593.9 608.4 690.3 779.9 817.0
936.0 499.0 443.0 1201.0 425.0 362.0 377.0 1226.0 1173.0 610.0 765.0 1170.0 458.0 321.0 525.0 322.0 379.0 1209.0 1527.0 457.0 1079.0 1193.0 333.0 439.0
329.8 456.9 484.0 1210.0 500.0 470.0 471.0 1188.0 1188.0 473.0 1134.0 1161.0 462.9 464.9 473.0 474.1 480.4 1189.0 1498.0 454.4 973.4 1158.0 427.5 442.4
606.1 638.5 755.1 660.7 700.2 829.3 773.9 804.0 888.2 937.9 1026.1 1219.2 Ages (Ma) 206 207 Pb Pb 238 206 U Pb
207
Pb 238 U
206
0.82 0.26 0.42
0.69 0.71 0.39 0.49 0.63 0.42 0.28 0.68 0.35 0.43 0.66 0.81 0.57 0.80
0.28 0.26 0.83 0.90 0.83 0.66 0.58
rho§§
0.93 0.92 0.96 0.91
0.96 0.83 0.53 0.92 0.98 0.70 0.44 0.98 0.78 0.56 0.75 0.75 0.86 0.83
rho§§
Ordovician calc-alkaline granitoids in the Acatlán Complex 469
470
Miller et al.
0.076
Palo Liso Granite 461.0 ± 2.1 Ma and 4 ± 53 Ma, MSWD = 0.54 (PL-1) ID-TIMS
U
0.074
A
462 458
238
454 850
450 0.072
446
206
Pb/
470 466
750
442 438 0.070
Intercepts: ca. 479 Ma & ca. 1065 Ma
650
550
0.09
0.068 0.53
0.54
207
Pb/
450
235
U
0.07 0.5
0.7
0.090
Palo Liso Granite
540
(PL-1) SHRIMP
B
0.086
0.082
500 0.078
1300
480
206
Pb/
238
U
520
1100
40
AR/ 39Ar GEOCHRONOLOGY
460
0.074
900
440
700
Analytical Methods
0.070 500
0.08
420
300
0.066 0.50
0.54
207
Pb/
235
U
0.00 0.0
0.4
0.8
0.078
238
238
Pb/ U weighted mean (LN-1) SHRIMP
Pb/
206
C
510 490 470 450 430
480 1400
Pb/
238
U
0.082
206
La Noria Granite 467 ± 16
U Age
0.086
206
an analysis at ca. 1210 Ma (Fig. 3B). A clearly distinguishable trend of four slightly discordant single-grain ID-TIMS analyses anchored by one concordant analysis (Fig. 3A) has an upper intercept of 461.0 ± 2.1 Ma (Fig. 3A), which is consistent with the youngest group of SHRIMP analyses (Fig. 3B). We consider this age to represent the time of pluton crystallization. Four ID-TIMS analyses of single zircon grains from the La Noria pluton also contain Mesoproterozoic inheritance (not plotted in Fig. 3; see Table 1) and do not provide any age interpretation. Seven SHRIMP analyses cluster around ca. 450–480 Ma, with a weighted mean of the 206Pb/238U ages of 467 ± 16 Ma, which we consider the best indication from these data for the time of pluton crystallization; other SHRIMP analyses are consistent with Mesoproterozoic inheritance (Fig. 3C and Table 1). The age of this pluton will be refined by additional ID-TIMS work. SHRIMP data for the Los Hornos pluton (Table 1) yielded one concordant point at ca. 442 Ma, one reversely discordant point with a 206Pb/238U age of ca. 428 Ma, and one slightly discordant point with a 207Pb/206Pb age of ca. 1200 Ma. Although these data are insufficient to provide an age interpretation, the youngest analyses are taken as an indication that the Los Hornos granite is probably part of this same general suite of Ordovician plutons and can be related temporally and petrogenetically to the other two granites. The age of this pluton will be refined by ID-TIMS analyses.
460
0.074
1000
440 0.070
420 600
0.10
0.066 0.50
0.54
207
Pb/
235
U
0.06 0
1
Figure 3. U-Pb Concordia diagrams for the (A and B) Palo Liso and (C) La Noria granites. MSWD—mean square of weighted deviates; ID-TIMS—isotope dilution thermal ionization mass spectrometry; SHRIMP—sensitive high-resolution ion microprobe.
Biotite and hornblende separates from the three megacrystic granitoid samples were pretreated and concentrated using standard techniques and later selected by hand-picking under a binocular microscope from fractions that ranged in size from 40 to 60 mesh in the mineral separation laboratory at Universidad Nacional Autónoma de México, Campus Juriquilla, Querétaro. Mineral separates were loaded into aluminum foil packets and irradiated together with Hb3gr (1072 Ma) as a neutron-fluence monitor at the McMaster Nuclear Reactor, Hamilton, Ontario. 40 Ar/39Ar analyses were performed by standard laser step-heating techniques described in detail by Clark et al. (1998) at the Geochronology Research Laboratory of Queens University, Kingston, Ontario. The data are given in Table 2 and plotted in Figure 4. All data have been corrected for blanks, mass discrimination, and neutron-induced interferences. For the purposes of this article, a plateau age is obtained when the apparent ages of at least three consecutive steps, comprising a minimum of 55% of the 39Ark released, agree within 2σ error with the integrated age of the plateau segment. The errors shown in Table 2 and on the age spectrum represent the analytical precision at ± 2σ. Results Hornblende from the La Noria pluton yielded a disturbed spectrum that monotonically decreases from ca. 429 Ma at the
Ar/40Ar
36
Ar/40Ar
39
r
TABLE 2. AR/AR ISOTOPIC DATA FOR ACATLÁN COMPLEX ORDOVICIAN GRANITES 39 40 40 Ca/K %40Atm %39Ar Ar*/39K Age Ar Ar Ar
38
Ar
37
0.022 ± 0.003 0.031 ± 0.003 0.022 ± 0.003 0.026 ± 0.003
0.013 ± 0.002 0.017 ± 0.002 0.010 ± 0.002 0.076 ± 0.003
0.064 ± 0.002 0.047 ± 0.003 0.103 ± 0.003 0.262 ± 0.004 0.315 ± 0.004 0.169 ± 0.003 0.508 ± 0.005
0.061 ± 0.002 0.460 ± 0.005 0.895 ± 0.007 1.068 ± 0.008
Notes: Atmospheric 40Ar/36Ar = 288.024. Measured volumes are × 1E-10 cm3 normal temperature and pressure (NTP). All errors are 2 × standard error. MSWD—mean square of weighted deviates.
Los Hornos granite: Biotite AOR R289 (LH-1; 18°24', 98°11'30'') 0.50 0.001684 ± 0.000375 0.059641 ± 0.000764 0.020 0.265 49.67 23.47 8.425 ± 1.863 112.69 ± 24.15 5.434 ± 0.040 0.331 ± 0.003 0.014 ± 0.003 <1.00> 0.000910 ± 0.000186 0.060136 ± 0.000665 0.033 0.233 26.84 38.24 12.157 ± 0.928 160.45 ± 11.72 8.767 ± 0.074 0.536 ± 0.004 0.016 ± 0.003 <2.00> 0.002295 ± 0.000188 0.028015 ± 0.000395 0.047 0.216 67.77 19.56 11.484 ± 1.998 151.93 ± 25.35 9.612 ± 0.068 0.278 ± 0.003 0.014 ± 0.003 <7.00> 0.002885 ± 0.000104 0.012564 ± 0.000189 0.107 2.308 85.22 18.73 11.739 ± 2.470 155.16 ± 31.28 20.483 ± 0.141 0.267 ± 0.003 0.030 ± 0.003 39 3 J value: 0.007651 ± 0.000058. Mass: 55.0 mg. Volume K: 1.37 × 1E-10 cm NTP. Approx. 0.13% Ca, 0.08% K. Integrated age: 146.68 ± 10.61 Ma. Initial 40/36: 296.56 ± 136.85 (MSWD = 1.91, isochron between –0.41 and 3.83). Correlation age: 152.13 ± 42.40 Ma (76.5% of 39Ar, steps marked by >). MSWD: 0.213. Plateau age: 156.98 ± 11.69 Ma (76.5% of 39Ar, steps marked by <). Mod. err.: 21.85.
La Noria granite: Biotite AOR-292 (LN-1; 18°24', 98°11'30'') 1.00 0.002555 ± 0.000140 0.037364 ± 0.000489 0.140 1.003 75.40 9.76 6.560 ± 1.120 88.06 ± 14.68 13.315 ± 0.132 0.507 ± 0.004 0.030 ± 0.003 2.00> 0.000145 ± 0.000080 0.041990 ± 0.000950 0.037 0.338 4.27 20.50 22.796 ± 0.778 289.13 ± 9.11 24.878 ± 0.530 1.061 ± 0.008 0.028 ± 0.004 3.00 0.000029 ± 0.000086 0.037181 ± 0.000739 0.007 1.021 0.84 15.60 26.668 ± 0.865 333.95 ± 9.89 21.372 ± 0.391 0.808 ± 0.006 0.020 ± 0.003 4.00 0.000036 ± 0.000051 0.033879 ± 0.000341 0.007 1.642 1.07 25.21 29.200 ± 0.533 362.65 ± 6.00 37.856 ± 0.321 1.298 ± 0.007 0.030 ± 0.003 5.00> 0.000029 ± 0.000075 0.032168 ± 0.000436 0.008 3.688 0.85 13.53 30.821 ± 0.805 380.81 ± 8.96 21.399 ± 0.260 0.701 ± 0.004 0.018 ± 0.002 <6.00> 0.000186 ± 0.000096 0.028393 ± 0.000934 0.016 5.563 5.49 4.78 33.284 ± 1.480 408.04 ± 16.23 8.562 ± 0.101 0.253 ± 0.007 0.020 ± 0.005 <7.00 0.000050 ± 0.000094 0.028861 ± 0.000300 0.013 7.609 1.48 10.62 34.135 ± 1.024 417.35 ± 11.18 18.691 ± 0.124 0.553 ± 0.004 0.020 ± 0.003 39 3 J value: 0.007625 ± 0.000064. Volume K: 5.08 × 1E-10 cm NTP. Approx. 5.15% K, 11.88% Ca. Integrated age: 328.78 ± 4.45 Ma. Initial 40/36: 738.84 ± 6966.04 (MSWD = 130.64, isochron between –0.41 and 3.83). Correlation age: 375.60 ± 267.50 Ma (38.8% of 39Ar, steps marked by >). MSWD: 0.893. Plateau age: 414.47 ± 9.74 Ma (15.4% of 39Ar, steps marked by <). Mod. err.: 9.74.
La Noria granite: Hornblende AOR-R293 (LN-1; 18°24’, 98°11’30’’) 2.00 0.000755 ± 0.000109 0.038819 ± 0.000413 0.009 0.729 22.28 22.98 20.016 ± 0.862 253.77 ± 10.19 16.724 ± 0.056 0.662 ± 0.007 4.00> 0.000045 ± 0.000062 0.037241 ± 0.000247 0.006 3.194 1.33 41.20 26.494 ± 0.523 328.81 ± 5.93 31.202 ± 0.069 1.178 ± 0.007 6.00> 0.000087 ± 0.000086 0.033487 ± 0.000317 0.020 10.801 2.57 23.86 29.093 ± 0.805 358.06 ± 8.98 20.039 ± 0.136 0.686 ± 0.004 <7.00> 0.000193 ± 0.000139 0.026493 ± 0.000355 0.052 25.906 5.70 11.95 35.595 ± 1.581 429.21 ± 16.97 12.618 ± 0.027 0.348 ± 0.004 J value: 0.007546 ± 0.000042. Mass: 8.0 mg. Volume 39K: 2.82 × 1E-10 cm3 NTP. Approx. 1.80% K, 12.82% Ca. Integrated age: 331.15 ± 4.82 Ma. Initial 40/36: 1680.98 ± 5625.21 (MSWD = 0.19, isochron between –0.41 and 3.83). Correlation age: 316.61 ± 194.39 Ma (77.0% of 39Ar, steps marked by >). Plateau age: 429.21 ± 17.10 Ma (12.0% of 39Ar, steps marked by <).
Palo Liso granite: Biotite AOR-R294 (PL-1; 18°30.667’, 98°17.778’’) 1.00 0.004361 ± 0.000718 0.074809 ± 0.003497 0.051 6.458 128.07 33.95 –3.791 ± 2.709 –52.34 ± 37.94 3.704 ± 0.016 0.286 ± 0.013 0.417 ± 0.379 0.221 ± 0.047 <2.00> 0.000326 ± 0.000297 0.039536 ± 0.000754 0.003 1.815 9.61 27.53 22.861 ± 2.263 286.97 ± 26.26 5.679 ± 0.022 0.244 ± 0.004 0.026 ± 0.003 0.055 ± 0.002 <3.00 0.000089 ± 0.000424 0.037814 ± 0.000891 0.003 1.697 2.62 18.48 25.752 ± 3.366 320.22 ± 38.35 3.991 ± 0.016 0.165 ± 0.003 0.014 ± 0.002 0.035 ± 0.002 <4.00 0.000043 ± 0.000936 0.034650 ± 0.001441 0.003 6.369 1.24 8.88 28.502 ± 8.060 351.28 ± 90.27 2.100 ± 0.012 0.085 ± 0.003 0.013 ± 0.003 0.060 ± 0.002 <6.00> 0.000863 ± 0.004178 0.035773 ± 0.005655 0.003 5.692 25.47 2.04 20.828 ± 34.658 263.22 ± 407.55 0.483 ± 0.007 0.022 ± 0.003 0.003 ± 0.003 0.013 ± 0.002 <7.00> 0.000184 ± 0.001023 0.040377 ± 0.001676 0.006 20.193 5.37 9.12 23.435 ± 7.540 293.62 ± 87.18 1.841 ± 0.028 0.079 ± 0.003 0.025 ± 0.003 0.183 ± 0.003 J value: 0.007542 ± 0.000042. Mass: 5.0 mg. Volume 39K: 0.81 × 1E-10 cm3 at normal temperature and pressure (NTP). Approx. 0.83% K, 57% Ca. Integrated age: 191.65 ± 21.64 Ma. 39 Correlation age: 300.40 ± 1948.98 Ma (38.7% of Ar, steps marked by >). MSWD: 0.86. Plateau age: 305.26 ± 26.17 Ma (66.0% of 39Ar, steps marked by <). Mod. err.: 22.05.
Power
0.011 ± 0.002 0.010 ± 0.002 0.025 ± 0.002 0.063 ± 0.002
0.038 ± 0.002 0.007 ± 0.002 0.004 ± 0.002 0.004 ± 0.002 0.004 ± 0.002 0.004 ± 0.001 0.004 ± 0.002
0.016 ± 0.002 0.005 ± 0.002 0.006 ± 0.002 0.006 ± 0.002
0.019 ± 0.003 0.005 ± 0.002 0.004 ± 0.002 0.003 ± 0.002 0.002 ± 0.002 0.002 ± 0.002
Ar
36
0.021 0.021 0.021 0.025
0.023 0.035 0.029 0.024 0.024 0.023 0.024
0.024 0.028 0.023 0.023
0.023 0.029 0.025 0.025 0.021 0.022
Blank 40 Ar
Ordovician calc-alkaline granitoids in the Acatlán Complex 471
472
Miller et al. 100
600
A
K/Ca
LN-1A (AOR-R293) La Noria
C
PL-1A (AOR-R294) Palo Liso
550
500
10
450
400
1
Hornblende
450
300
250
350
Age (Ma)
Age (Ma)
400
350
LN-1A (AOR-R293) La Noria
300
250
Hornblende
305 ± 26Ma 200
150
450
400
100
LN-1A (AOR-292) La Noria
B
Biotite
50
350 200
Age (Ma)
300
LH-1A (AOR-R289)
150 250 100 200
157 ± 12Ma
50 150
Los Hornos 0 0.0
100
Biotite F raction 39A r
D 1.0
50
Biotite 0 0.0
F raction 39A r
1.0
Figure 4. 40Ar/39Ar incremental release spectra from (A) hornblende and (B) biotite in the La Noria pluton; (C) biotite in the Palo Liso pluton; (D) biotite in the Los Hornos pluton.
highest temperature increment to ca. 254 Ma at the lowest power increment (Fig. 4A). Biotite from the same sample yielded a similar disturbed spectrum, stepping down from ca. 417 ± 11 Ma to ca. 88 Ma (Fig. 4B). Biotite from the Palo Liso pluton also yielded a disturbed spectrum that yielded a plateau age of 305 ± 26 Ma in the five highest power increments (Fig. 4C). Biotite from the Los Hornos pluton yielded a plateau age of 157 ± 12 Ma in the three highest power increments, stepping down to ca. 113 Ma in the lowest increment (Fig. 4D). Secondary alteration of the granitoids has clearly affected these rocks; epidote-rich veins are locally abundant, and in many samples biotite is partly altered to chlorite. Although an effort
was made to sample the most pristine area of an outcrop, all four spectra show significant disturbances in their 40Ar/39Ar systematics, which may be due to secondary alteration. Thus only the plateau ages from the Palo Liso and Los Hornos plutons are likely to be geologically meaningful, representing the time of cooling through the roughly 300 °C closure temperature of biotite (Harrison et al., 1985). The ca. 254 Ma thermal overprint may be associated with the widespread Permo-Triassic tectonothermal event recorded elsewhere in the Acatlán Complex (Keppie et al., 2004a, 2006), and the ca. 157 Ma overprint may be a localized result of the Jurassic migmatization event recorded in the eastern part of the Acatlán Complex (Keppie et al., 2004b).
Ordovician calc-alkaline granitoids in the Acatlán Complex
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SiO2 (wt%) TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5 LOI Total
TABLE 3. MAJOR AND TRACE ELEMENT COMPOSITION OF ACATLÁN COMPLEX ORDOVICIAN GRANITES LH-2 LH-3 LH-4 LN-2 LN-3 LN-5 PL-10 PL-11 PL-12 PL-13 71.46 70.02 70.46 69.18 68.53 65.56 74.61 71.93 72.67 71.27 0.55 0.63 0.50 0.75 0.80 0.55 0.21 0.44 0.54 0.50 13.72 13.67 14.06 14.41 14.45 15.65 12.52 13.42 13.52 13.16 3.69 4.24 3.44 5.03 5.16 4.30 2.08 3.21 1.87 3.60 0.05 0.06 0.05 0.09 0.07 0.06 0.03 0.06 0.05 0.09 0.98 1.11 0.89 1.44 1.63 1.75 0.55 0.86 1.61 1.01 0.78 0.89 1.01 2.36 1.57 2.85 0.20 0.64 0.52 1.76 2.82 2.54 2.66 2.96 2.41 3.34 2.21 2.67 3.46 2.61 4.79 4.51 5.42 2.73 4.39 3.20 5.74 4.95 3.20 4.63 0.13 0.14 0.13 0.16 0.16 0.12 0.12 0.12 0.14 0.13 1.77 1.80 1.68 1.46 1.54 1.98 0.82 1.63 1.55 2.50 100.74 99.61 100.30 100.56 100.71 99.37 99.09 99.94 99.12 101.26
PL-14 60.26 1.50 15.28 9.81 0.16 2.73 1.71 2.39 4.24 0.23 2.35 100.66
Cr (ppm) Co V Pb Zn Rb Ba Sr Ga Nb Zr Y Th La Nd
16 11 64 18 40 158 757 67 17 13 237 35 12 39 28
38 30 163 13 124 146 468 115 25 33 574 51 22 83 55
21 12 67 14 59 168 495 66 18 18 257 39 12 52 35
14 7 59 19 48 167 1323 86 15 12 200 34 12 65 53
14 15 81 13 64 79 555 171 18 16 313 43 16 64 45
22 13 95 18 73 153 690 118 17 16 305 41 17 72 50
GEOCHEMISTRY Analytical Methods Eleven samples of megacrystic granite were collected for chemical analyses: five from the Palo Liso pluton, three from the Los Hornos pluton, and three from the La Noria pluton. The samples were analyzed by X-ray fluorescence spectrometer for major elements and several trace elements (Rb, Sr, Ba, Zr, Nb, Y, Zn, V, Cr, and Ni) in the Nova Scotia Regional Geochemical Centre at Saint Mary’s University, Halifax. The precision and accuracy of the X-ray data are reported by Dostal et al. (1994). The analyses of representative samples are given in Table 3. Whole-Rock Geochemistry The rocks of the Los Hornos pluton are granitic in composition, with SiO2 contents of ~70 wt% accompanied by high alkalis and K2O/Na2O ratios (~1.7–2.0). The rocks of the La Noria pluton range in SiO2 between 65 and 70 wt%, but have slightly lower total alkalis, with K2O/Na2O between 0.9 and 1.8. The rocks of the Palo Liso pluton have a wider SiO2 range (60–75 wt%) although most rocks contain 71–75 wt% SiO2. They are high in alkalis (6.8–8.1 wt%) and typically have high K2O/Na2O ratios (1.0–2.6). According to classification of Middlemost (1985), based on the (Na2O + K2O) versus SiO2 diagram (Fig. 5A), the rocks are granites and subordinate granodiorites, and one sample corresponds to quartz monzodiorite. Likewise, on the Q-A-P plot
1 12 63 21 42 110 550 263 18 13 244 26 12 46 37
6 0 23 6 18 229 745 27 11 11 91 23 12 22 18
19 8 45 13 46 174 561 69 16 13 172 32 9 40 31
18 0 61 0 12 126 1152 88 16 15 215 33 8 36 29
17 10 53 11 50 163 463 56 16 12 205 34 11 51 36
of LeMaitre (1989), the rocks of all three groups fall predominantly into the field of monzogranite; the rest are granodiorites (Fig. 5B). The granitic rocks are peraluminous [molar Al2O3 > (CaO + Na2O + K2O)], with corundum in their norms and calcalkaline affinities. The granitic rocks of the three groups have overlapping major and trace element compositions (Table 3) and relatively uniform ratios of Zr/Y (4–10), Nb/Y (0.3–0.5), Ti/V (50–60), Ti/Zr (13–16), and K/Rb (200–300), suggesting that they have shared a similar origin and evolution. DISCUSSION AND CONCLUSIONS The geochronologic and geochemical data show that the three undeformed megacrystic plutons in the northern part of the Acatlán Complex are Middle Ordovician to early Silurian (ca. 470–440 Ma) and have calc-alkaline affinities. The ca. 1200 Ma inheritance ages recorded in the zircons from the three plutons are consistent with a source in or contamination from sediments derived from a Mesoproterozoic geological province. The most immediate candidate for such a source of Mesoproterozoic zircons is the Oaxacan Complex, which is inferred to underlie the Acatlán Complex (Keppie et al., 2003a; Solari et al., 2003; Keppie, 2004). This would be consistent with an Ordovician paleogeography that places the Acatlán Complex adjacent to Oaxaquia (Keppie, 2004) along the Gondwanan margin of the Rheic Ocean. The calc-alkaline geochemistry of these Ordovician granitoids is also consistent with melting of the Oaxaquia basement, part of which has a calcalkaline signature (Lawlor et al., 1999; Keppie et al., 2001).
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B
(wt %)
10
Na2O + K2O
A
Quartz monzonite
Q
Granite
La Noria Palo Liso Los Hornos Granodiorite
5 Tonolite
Los Hornos La Noria Palo Liso
Monzo granite
A
0 60
70
SiO2
(wt %)
Grano diorite
e o l it
Syeno granite
Ton
Quartz Monzodiorite
P
80
Figure 5. (A) Na2O + K2O versus SiO2 (wt.%) diagram of Middlemost (1985) for the Ordovician megacrystic granitoid rocks of the Acatlán Complex. (B) Quartz-alkali feldspar-plagioclase (QAP) plot for the normative composition of the Ordovician granitoid rocks of the Acatlán Complex. Fields after LeMaitre (1989).
The undeformed plutons in this study are lithologically similar to and about the same age as or slightly older than penetratively deformed megacrystic granitoid rocks elsewhere in the Acatlán Complex, which range from Tremadocian to Middle Ordovician (478 ± 5 Ma, Campa-Uranga et al., 2002; 471 ± 13 Ma, SánchezZavala et al., 2004). The deformed granitoids are inferred to have been located near the edge of the continental margin that was subsequently subducted, producing the Devono-Carboniferous eclogite-facies metamorphism (Middleton et al., this volume). A slightly later plutonic event in the early Silurian is recorded by the Xayacatlán metagabbro in the eastern part of the Acatlán Complex (Fig. 1). Most recent palinspastic reconstructions place Avalonia adjacent to Oaxaquia and Amazonia in the latest Precambrian (e.g., Keppie et al., 2003b, and references therein). Transtensional rifting may have resulted from collision of a mid-ocean ridge and a subduction zone, producing a Baja California–type margin. In this model, Avalonia would have separated from Gondwana, traveled along the northern Gondwanan margin during the Cambrian, and approached northern Gondwana during the Late Cambrian and Early Ordovician (Keppie et al., 2003b; Nance et al., this volume, Chapter 21). The megacrystic granites discussed here are inferred to be a consequence of ca. 480–440 Ma rift-related magmatism associated with the initial separation of Avalonia from Oaxaquia along the Gondwanan margin of the Rheic Ocean. Subsequent tectonothermal events recorded by the 40Ar/39Ar data in the three plutons studied here may be related to the Permo-Triassic and Jurassic events recorded elsewhere in the Acatlán Complex (Malone et al., 2002; Keppie et al., 2004a,b, 2006), and other lower-temperature, later disturbances to the 40Ar/39Ar data may be the result of the Laramide orogeny.
ACKNOWLEDGMENTS We would like to acknowledge PAPIIT (Programa de Apoyo a Proyectos de Investigación e Innovación Tecnológica) grant IN103003 to JDK, National Science Foundation grants EAR 0308105 to RDN and EAR 0456180 to BVM, a NSERC (National Science and Engineering Research Council Discovery) grant to JD and JKWL, and CONACyT (Consejo Nacional de Ciencia y Tecnología) grant 33100-T to AOR. We thank Jim Hibbard and Fritz Finger for constructive reviews and Miguel Morales for assistance with drawing the figures. This article represents a contribution to International Geological Correlation Program Project 498, The Rheic Ocean. REFERENCES CITED Campa, M.F., and Coney, P.J., 1983, Tectono-stratigraphic terranes and mineral distributions in Mexico: Canadian Journal of Earth Sciences, v. 20, p. 1040–1051. Campa-Uranga, M.F., Gehrels, G., and Torres de Leon, R., 2002, Nuevas edades de granitoides metamorfizados del complejo Acatlán en el Estado de Guerrero: Actas Instituto nacional de Geoquimica, v. 8, no. 1, p. 248. Clark, A.H., Archibald, D.A., Lee, A.W., Farrar, E., and Hodgson, C.J., 1998, Laser probe 40Ar/39Ar ages of early- and late-stage alteration assemblages, Rosario porphyry copper-molybdenum deposit, Collahuasi District, I Region, Chile: Economic Geology and the Bulletin of the Society of Economic Geologists, v. 93, p. 326–337. DeGraaff-Surpless, K., Graham, S.A., Wooden, J.L., and McWilliams, M.O., 2002, Detrital zircon provenance analysis of the Great Valley Group, California: Evolution of an arc-forearc system: Geological Society of America Bulletin, v. 114, p. 1564–1580, doi: 10.1130/0016-7606(2002)114<1564: DZPAOT>2.0.CO;2. Derycke-Khatir, C., Vachard, D., Dégardin, J.-M., Flores de Dios, A., Buitrón, B., and Hansen, M., 2005, Late Pennsylvanian and Early Permian chondichthyan microremains from San Salvador Patlanoaya (Puebla, Mexico): Geobios, v. 38, p. 43–55, doi: 10.1016/j.geobios.2003.06.008. Dostal, J., Dupuy, C., and Caby, R., 1994, Geochemistry of the Neoproterozoic Tilemsi belt of Iforas (Mali, Sahara), a crustal section of an oceanic
Ordovician calc-alkaline granitoids in the Acatlán Complex island arc: Precambrian Geology, v. 65, p. 55–69, doi: 10.1016/03019268(94)90099-X. Fries, C., and Rincon-Orta, C., 1965, Nuevas aportaciones geochronologicas y tecnicas empleadas en el laboratorio: Universidad Nacional Autónoma de México, Boletín del Instituto Geologico, v. 73, p. 57–133. Halpern, M., Guerrero-Garcia, J.C., and Ruiz-Castellanos, M., 1974, Rb-Sr dates of igneous and metamorphic rocks from southeastern and central Mexico: A progress report: Union Geofisica Mexicana, Reunion Annual, Resumenes, p. 30–31. Harrison, T.M., Duncan, I., and McDougall, I., 1985, Diffusion of 40Ar in biotite: Temperature, pressure, and compositional effects: Geochimica et Cosmochimica Acta, v. 49, p. 2461–2468. Keppie, J.D., 2004, Terranes of Mexico revisited: A 1.3 billion year odyssey: International Geology Review, v. 46, p. 765–794. Keppie, J.D., Dostal, J., Ortega-Gutiérrez, F., and Lopez, R., 2001, A Grenvillian arc on the margin of Amazonia: Evidence from the southern Oaxacan Complex, southern Mexico: Precambrian Research, v. 112, p. 165–181, doi: 10.1016/S0301-9268(00)00150-9. Keppie, J.D., Dostal, J., Cameron, K.L., Solari, L.A., Ortega-Gutiérrez, F., and Lopez, R., 2003a, Geochronology and geochemistry of Grenvillian igneous suites in the northern Oaxacan Complex, southern México: Tectonic implications: Precambrian Research, v. 120, p. 365–389, doi: 10.1016/ S0301-9268(02)00166-3. Keppie, J.D., Nance, R.D., Murphy, J.B., and Dostal, J., 2003b, Tethyan, Mediterranean, and Pacific analogues for the Neoproterozoic–Paleozoic birth and development of peri-Gondwanan terranes and their transfer to Laurentia and Laurussia: Tectonophysics, v. 365, no. 1–4, p. 195–219, doi: 10.1016/S0040-1951(03)00037-4. Keppie, J.D., Sandberg, C.A., Miller, B.V., Sánchez-Zavala, J.L., Nance, R.D., and Poole, F.G., 2004a, Implications of latest Pennsylvanian to Middle Permian paleontological and U-Pb SHRIMP data from the Tecomate Formation to re-dating tectonothermal events in the Acatlán Complex, southern Mexico: International Geology Review, v. 46, p. 745–754. Keppie, J.D., Nance, R.D., Powell, J.T., Mumma, S.A., Dostal, J., Fox, D., Muise, J., Ortega-Rivera, A., Miller, B.V., and Lee, J.W.K., 2004b, MidJurassic tectonothermal event superposed on a Paleozoic geological record in the Acatlán Complex of southern Mexico: Hotspot activity during the breakup of Pangea: Gondwana Research, v. 7, no. 1, p. 239–260. Keppie, J.D., Nance, R.D., Fernandez-Surez, J., Storey, C.D., Jeffries, T.E., and Murphy, J.B., 2006, Detrital zircon data from the eastern Mixteca terrane, southern Mexico: Evidence for an Ordovician–Mississippian continental rise and a Permo-Triassic clastic wedge adjacent to Oaxaquia: International Geology Review, v. 48, p. 97–111. Lawlor, P.J., Ortega-Gutiérrez, F., Cameron, K.L., Ochoa-Camarillo, H., Lopez, R., and Sampson, D.E., 1999, U–Pb geochronology, geochemistry, and provenance of the Grenvillian Huiznopala Gneiss of Eastern Mexico: Precambrian Research, v. 94, p. 73–99, doi: 10.1016/S0301-9268(98)00108-9. LeMaitre, R.W., 1989, A classification of igneous rocks and a glossary of terms: Oxford, England, Blackwell, 193 p. Ludwig, K.R., 1989, Pb-Dat: A computer program for processing raw Pb-U-Th isotope data: Reston, Virginia, U.S. Geological Survey Open File Report 88-557, 34 p. Ludwig, K.R., 2001, SQUID 1.02: A user’s manual: Berkeley, California, Berkeley Geochronology Center Special Publication 2. Ludwig, K. R., 2003, Isoplot/Ex, version 2.49: A geochronological toolkit for Microsoft Excel: Berkeley, California, Berkeley Geochronology Center Special Publication 1a. Malone, J.W., Nance, R.D., Keppie, J.D., and Dostal, J., 2002, Deformational history of part of the Acatlán Complex: Late Ordovician–Early Silurian and Early Permian orogenesis in southern Mexico: Journal of South American Earth Sciences, v. 15, p. 511–524, doi: 10.1016/S08959811(02)00080-9. Middlemost, E.A.K., 1985, Magmas and magmatic rocks: An introduction to igneous petrology: London, Longman, 266 p. Middleton, M., Keppie, J.D., Murphy, J.B., Miller, B.V., Nance, R.D., OrtegaRivera, A., and Le, J.K.W., 2007 (this volume), P-T-t constraints on
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exhumation following subduction in the Rheic Ocean from eclogitic rocks in the Acatlán Complex of southern México, 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: Boulder, Colorado, Geological Society of America Special Paper 423, doi: 10.1130/2007.2423(25). Miller, B.V., Fetter, A.H., and Stewart, K.G., 2006, Plutonism in three orogenic pulses, Eastern Blue Ridge Province, southern Appalachians: Geological Society of America Bulletin, v. 118, p. 171–184, doi: 10.1130/B25580.1. Nance, R.D., Miller, B.V., Keppie, J.D., Murphy, J.B., and Dostal, J., 2007 (this volume), Vestige of the Rheic Ocean in North America: The Acatlán Complex of southern México, 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: Boulder, Colorado, Geological Society of America Special Paper 423, doi: 10.1130/2007.2423(21). Ortega-Gutiérrez, F., 1978, Estratigrafia del Complejo Acatlán en la Mixteca Baja, Estados de Puebla y Oaxaca: Universidad Nacional Autónoma de México, Instituto de Geología, Revista, v. 2, p. 112–131. Ortega-Gutiérrez, F., 1981, Metamorphic belts of southern Mexico and their tectonic significance: Geofisica Internacional, v. 20, p. 177–202. Ortega-Gutiérrez, F., Elías-Herrera, M., Reyes-Salas, M., Macias-Romo, C., and López, R., 1999, Late Ordovician–Early Silurian continental collision orogeny in southern Mexico and its bearing on Gondwana-Laurentia connections: Geology, v. 27, p. 719–722, doi: 10.1130/00917613(1999)027<0719:LOESCC>2.3.CO;2. Ramirez-Espinosa, J., 2001, Tectono-magmatic evolution of the Paleozoic Acatlán Complex in southern Mexico, and its correlation with the Appalachian system [Ph.D. thesis]: Tucson, University of Arizona, 170 p. Reyes-Salas, A.M., 2003, Mineralogia y petrologia de los granitoides Esperanza del complejo Acatlán, sur de México [doctoral thesis]: Universidad Nacional Autónoma de México, 165 p. Robison, R., and Pantoja-Alor, J., 1968, Tremadocian trilobites from Nochixtlan region, Oaxaca, Mexico: Journal of Paleontology, v. 42, p. 767–800. Sánchez-Zavala, J.L., Ortega-Gutiérrez, F., and Elías-Herrera, M., 2000, La orogenia Mixteca del Devónico del complejo Acatlán, sur de Mexico: Geos Unión Geofísica Mexicana, Boletin Informativo Época II, v. 20, no. 3, p. 321–322. Sánchez-Zavala, J.L., Ortega-Gutiérrez, F., Keppie, J.D., Jenner, G.A., Belousova, E., and Maciás-Romo, C., 2004, Ordovician and Mesoproterozoic zircons from the Tecomate Formation and Esperanza granitoids, Acatlán Complex, southern Mexico: Local provenance in the Acatlán and Oaxacan complexes: International Geology Review, v. 46, p. 1005–1021. Solari, L.A., Keppie, J.D., Ortega-Gutiérrez, F., Cameron, K.L., Lopez, R., and Hames, W.E., 2003, 990 Ma and 1,100 Ma Grenvillian tectonothermal events in the northern Oaxacan Complex, southern Mexico: Roots of an orogen: Tectonophysics, v. 365, p. 257–282, doi: 10.1016/S00401951(03)00025-8. Talavera-Mendoza, O., Ruíz, J., Gehrels, G.E., Meza-Figueroa, D.M., VegaGranillo, R., and Campa-Uranga, M.F., 2005, U-Pb geochronology of the Acatlán Complex and implications for the Paleozoic paleogeography and tectonic evolution of southern Mexico: Earth and Planetary Science Letters, v. 235, p. 682–699, doi: 10.1016/j.epsl.2005.04.013. Vachard, D., Flores de Dios, A., Buitron, B.E., and Grahales, M., 2000, Biostratigraphie par fusilines des calcaires Carbonifères et Permienes de San Salvador Patlanoaya (Puebla, Mexique): Geobios, v. 33, p. 5–33, doi: 10.1016/S0016-6995(00)80145-X. Yañez, P., Ruiz, J., Patchett, P.J., Ortega-Gutiérrez, F., and Gehrels, G., 1991, Isotopic studies of the Acatlán Complex, southern Mexico: Implications for Paleozoic North American tectonics: Geological Society of America Bulletin, v. 103, p. 817–828, doi: 10.1130/0016-7606(1991)103<0817: ISOTAC>2.3.CO;2.
MANUSCRIPT ACCEPTED BY THE SOCIETY 3 OCTOBER 2006
Printed in the USA
Geological Society of America Special Paper 423 2007
Ordovician–Devonian oceanic basalts in the Cosoltepec Formation, Acatlán Complex, southern México: Vestiges of the Rheic Ocean? J. Duncan Keppie* Departamento de Geología Regional, Instituto de Geología, Universidad Nacional Autónoma de México, 04510 México D.F. Jaroslav Dostal Department of Geology, St. Mary’s University, Halifax, Nova Scotia, Canada B3H 3C3 Mariano Elías-Herrera Departamento de Geología Regional, Instituto de Geología, Universidad Nacional Autónoma de México, 04510 México D.F. ABSTRACT The Cosoltepec Formation consists of unfossiliferous phyllites and psammites with tectonic slices of basalt and forms a major part of the Acatlán Complex of southern México. It has traditionally been interpreted as part of either a Cambro-Ordovician accretionary prism or a passive margin deposited in the Iapetus Ocean. Although no reliable age data are available from the basaltic slices, their widespread tectonic interleaving with the Cosoltepec Formation suggests that some of the sedimentary rocks were originally deposited directly on the ocean-floor basalts. The age of the Cosoltepec metasediments ranges from Ordovician to uppermost Devonian (oldest unconformably overlying sediments: uppermost Devonian). The mafic rocks have been affected by greenschist- to sub-greenschist-facies metamorphism. The geochemistry indicates that they are mainly mid-ocean ridge basalt (MORB) tholeiites with flat or depleted rare earth element (REE) patterns and resemble MORB derived from heterogeneous sources. Another minor group consists of ocean island basalts and andesites, which have distinctly fractionated REE patterns, whereas their mantle-normalized trace element patterns do not show Nb-Ta or Ti negative anomalies. Thus, these mafic rocks appear to represent oceanic lavas that were tectonically incorporated into the clastic rocks of the Cosoltepec Formation during deformation that started immediately prior to deposition of late Fammenian sedimentary rocks and continued into the Mississippian. Tectonic juxtaposition of these rocks with eclogitic rocks suggests that the deformation was related to exhumation following subduction: the Cosoltepec Formation and its mafic lenses derived from the overriding plate, whereas the eclogitic rocks represent parts of the subducting plate. The probable Middle Ordovician–Middle Devonian age of the oceanic rocks together with their initial deformation during the Devonian–Carboniferous suggest that they represent vestiges of the Rheic Ocean. Keywords: Acatlán, México, oceanic basalt, Rheic *E-mail:
[email protected]. Keppie, J.D., Dostal, J., and Elías-Herrera, M., 2007, Ordovician–Devonian oceanic basalts in the Cosoltepec Formation, Acatlán Complex, southern México: Vestiges of the Rheic Ocean?, 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. 477–487, doi: 10.1130/2007.2423(24). For permission to copy, contact
[email protected]. ©2007 Geological Society of America. All rights reserved.
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INTRODUCTION The Acatlán Complex of southern México (Fig. 1) is inferred to be the vestige of a Paleozoic ocean that has been correlated with either the Iapetus Ocean (Ortega-Gutiérrez et al., 1999), or the Rheic Ocean (Keppie and Ramos, 1999). It is separated on its eastern side from the ca. 1 Ga Oaxacan Complex by a Permian dextral fault zone (Elías-Herrera and Ortega-Gutiérrez, 2002). The Oaxacan Complex is part of the Oaxaquia terrane, which is inferred to have lain adjacent to the Amazon craton throughout the Paleozoic (Keppie, 2004). The presence of oceanic rocks in the Acatlán Complex is based primarily on the existence of pillow basalts in the Cosoltepec Formation; however, geochemical data are almost nonexistent (cf. Ramírez-Espinosa, 2001). In order to fill this gap, we present geochemical data from the mafic rocks of the Cosoltepec Formation, which document that they are predominantly ocean floor tholeiites (MORB) with minor ocean island basalts (OIB). GEOLOGICAL SETTING Until recently, the Acatlán Complex, which constitutes the Mixteca terrane of southern México, was inferred to have undergone the following sequence of events (Fig. 2, left-hand side): (1) Cambro-Ordovician deposition of the clastic Petlalcingo Group (the Magdalena, Chazumba, and Cosoltepec Formations) and the oceanic Piaxtla Group (the Xayacatlán Formation and Esperanza Granitoids); (2) Late Ordovician–Early Silurian polyphase deformation during the Acatecan orogeny, when the Piaxtla Group underwent eclogite-facies metamorphism and was thrust over the Petlalcingo Group metamorphosed at greenschist-amphibolite facies; (3) Devonian deposition of the Tecomate Formation; (4) Devonian deformation and greenschist-facies metamorphism during the Mixtecan orogeny; and (5) deposition of the uppermost Devonian–Lower Permian Patlanoaya Formation (Ortega-Gutiérrez et al., 1999; SánchezZavala et al., 2000; Vachard and Flores de Dios, 2002). However, recent work has shown significant changes in age assignments of the various units and tectonothermal events (Fig. 2, right-hand side). Thus, the ages of the youngest detrital zircons in the Magdalena and Chazumba Units indicate that they are Permo-Triassic (Keppie et al., 2006). The Cosoltepec Formation consists of extensive phyllites and psammites with tectonic slices of mafic metavolcanic rocks that rarely preserve pillows with intercalated red chert, e.g., in the southwestern part of the area (Fig. 1). To date, the red cherts have yielded only algae with a wide age range. Thus, although the age of the basaltic rocks is unknown, the ages of the enclosing metasedimentary rocks of the Cosoltepec Formation range from preMiddle or pre-Upper Ordovician to pre-uppermost Devonian, suggesting that it is a composite unit: (1) some metasedimentary rocks are intruded by Middle–Upper Ordovician granitoids (Miller et al., this volume), (2) the youngest detrital zircon from Cosoltepec metasedimentary rocks immediately east of Acatlán
yielded a concordant U-Pb age of ca. 455 Ma (Keppie et al., 2006), (3) metasedimentary rocks from the type area and the northernmost outcrops have yielded detrital zircons as young as ca. 410 Ma (Talavera-Mendoza et al., 2005): samples COS 9–22 come from the type area but are bounded by thrust faults (Fig. 1B), and (4) undeformed rocks of the uppermost Devonian–Lower Permian Patlanoaya Formation rest unconformably on various units of strongly deformed rocks of the Acatlán Complex (Vachard and Flores de Dios, 2002). Unfortunately, the mafic rocks sampled for this study are not associated with Cosoltepec metasedimentary rocks cut by the Ordovician granitoids. Furthermore, they are always in tectonic contact with the enclosing metasedimentary rocks, so detrital zircons cannot directly constrain their age. On the other hand, the undeformed nature of the Patlanoaya Formation compared with the intensely deformed Cosoltepec Formation would appear to provide a younger, latest Devonian, age constraint. The close association of clastic metasedimentary rocks and the mafic volcanic slices throughout the Acatlán Complex suggests that the metasediments were deposited on the basalts as a result of progradation of the continental sediments over the ocean floor (Keppie et al., 2006). An older age constraint is suggested by the fact that the oldest dated igneous rocks in the Acatlán Complex are ca. 440–480 Ma (Dostal et al., 2003; Keppie et al., 2004c; Miller et al., this volume) and that the oldest continental shelf rocks unconformably overlying the adjacent ca. 1 Ga Oaxacan Complex are latest Cambrian–earliest Ordovician (Tremadocian) (Landing et al., 2006). Fossils and dated pebbles in the Tecomate Formation indicate that it is uppermost Carboniferous–Middle Permian (Keppie et al., 2004a). The type-Xayacatlán igneous body has yielded a concordant U-Pb zircon age of 442 ± 2 Ma (lowest Silurian) and has a continental tholeiitic signature (Dostal et al., 2003). The tectonothermal events have been dated as (1) Late Devonian–Mississippian (eclogite-facies metamorphism; Middleton et al., this volume); (2) Permian–Triassic (greenschist-facies metamorphism; Malone et al., 2002; Keppie et al., 2004a, 2006); and (3) Jurassic (local migmatization and high-temperature / low-pressure metamorphism; Keppie et al., 2004b). The Cosoltepec Formation has been interpreted as an accretionary prism (Ortega-Gutiérrez et al., 1999), a passive margin sequence (Ramírez-Espinosa, 2001), or a continental rise deposit (Keppie et al., 2006). In order to resolve the tectonic setting in which the basaltic lavas were extruded, we sampled the basaltic rocks at various locations across the Acatlán Complex (Fig. 1). Three samples (IXCA 2, 4, and 5) collected in the high-grade Xayacatlán Formation in the western Mixteca terrane (Fig. 1) are tectonically interleaved with pelitic muscovite-biotite-chlorite schist. A chlorite-phengite sample from nearby has yielded detrital zircon ages with a young peak at ca. 477 Ma (Talavera-Mendoza et al., 2005), the latter providing an older age constraint on deposition of the metasedimentary rocks.
Ordovician–Devonian oceanic basalts in the Cosoltepec Formation
America
A
s to
aja
Cortez
Laramide front
Sier Mad r a re
Oaxaquia (Middle America)
Tecomate & Patlanoaya Fms. u Tro
GU
Zihuatenajo Maya Las Ollas RE (M.America) Arcelia TE RO Teloloapan ua RR C tag Mo AN OM Mixteca E PO SI TE Fig. 1B Oaxaquia 250 500 Km (Middle America)
n yma
Ca
Totoltepec pluton (288 Ma)
gh
Ordovician-Mississippian
Nicaragua Rise
ar Ju
ER
ez
0
Post-Permian cover Permo-Triassic Chazumba Petlalcingo Lithodeme Magdalena Suite Lithodeme Mississippian-Middle Permian
la
Cenozoic arc volcanic rocks
C oa hu i
isi Al
ino . B ca W Viosyal Ch
North
Ac a t lá n C o m p le x
Ouachita front
Tarahumara
479
Cosoltepec Fm.
Chortis (M.America) ent m arp sc E ss He
Esperanza Gt (Ordovician) Xayacatlán Fm. Ordovician
* Sample locations & numbers (all have prefix COS except IXCA)
Granitoid plutons Precambrian ca. 1 Ga Oaxacan Complex
B
Ri
ve
r
98° 30´ W
Piaxtla Suite
Atoy ac
190
Izúcar de Matamoros
Tehuacán Patlanoaya
r
ve
18° 30´ N
Ri a
p xa
3-8
*
Ne
IXCA 2,3,&5
15-19 20-22 Ac atlán
r
* * o Mixtec
9-14
Huajuapan
18° 00´ N
Rive r
39-45
co
* 46-54 * *
ane
Tlap
Hermanns, 1994 Olinalá
Caltepec
Caltepec fault zone
Ba
Riv e
ls
as
**
Metzontla
Tamazulapan
Rive
r
23-37
0
30 km 97° 30´ W
190
Figure 1. (A) Map of Middle America showing the location of the Mixteca terrane. (B) geological map of the Mixteca terrane (modified after Keppie et al., 2006) showing the locations of the mafic samples described in this article and the locations of samples in Hermanns’s (1994) thesis. Fm.—Formation; Gt—Granite.
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TIME SCALE (Ma) 200 ORTEA GUTIÉRREZ ET AL., 1999
P
*
300
350
400
450
500
Patlanoaya Formation
C D S O C
* Mixtecan Orogeny Tecomate Formation Acatecan Orogeny Cosoltepec Chazumba Magdalena IAPETUS OCEAN
550
Petlalcingo Group
250
Tr
THIS PAPER Green Chazumba schist Magdalena facies * (Petlalcingo event Suite) * * Tecomate and Patlanoaya Fms Otates Fm * Eclogite facies event Cosoltepec Formation Continental rise & oceanic lavas RHEIC OCEAN
Figure 2. Time and space diagram showing the geological history of the Mixteca terrane according to Ortega-Gutiérrez et al. (1999) and this article. Tr—Triassic; P—Permian; C— Carboniferous; D—Devonian; S—Silurian; O—Ordovician; –C—Cambrian.
455+/-4
* Fossil Youngest detrital zircon
PETROGRAPHY
GEOCHEMISTRY
The mafic volcanic assemblage within the Cosoltepec Formation has undergone regional greenschist and sub-greenschist-grade metamorphism. This has changed the primary mineralogy of the rocks, although relicts of igneous textures are commonly seen under the microscope. The mafic lavas are aphyric to porphyritic and are composed mainly of secondary phases. In thin section, the lavas consist of randomly oriented or subparallel laths of plagioclase (0.2–1.0 mm long) and, in the southwestern part of the area, clinopyroxene and olivine phenocrysts set in a groundmass composed of chlorite, albite, epidote, and Fe-Ti oxides with or without calcite, quartz, actinolite, and pyrite. Calcic plagioclase that has been largely transformed to albite and chlorite generally replaces the primary ferromagnesian minerals. Amygules are filled with quartz, calcite, chlorite, pumpellyite, and epidote. The two mafic samples (IXCP 2 and 5) collected from within the Xayacatlán Formation in the western part of the Mixteca terrane (Fig. 1) contain blue-green sodic amphibole, albite, epidote, chlorite, quartz, opaque minerals, and calcite. An adjacent outcrop consists of garnetiferous amphibolites: Talavera-Mendoza et al. (2005) recorded blueschists from nearby outcrops, suggesting that the amphibolites might be retrogressive after blueschist.
Analytical Methods Fifty-five samples of volcanic rocks of the Cosoltepec Formation and two samples of the Xayacatlán Formation have been analyzed for major and trace elements. Analyses of major elements and some trace elements (Rb, Sr, Ba, Zr, Nb, Y, La, Ce, Nd, Ga, Cu, Zn, V, Sc, Co, Cr, and Ni) were done using X-ray fluorescence (fused glass discs and powder pellets) at the Geochemical Centre of Saint Mary’s University, Halifax. The precision and accuracy of these analyses are discussed in Dostal et al. (1986, 1994). Nine representative samples were chosen for determination of a wider range of trace elements including rare earth elements (REE), Th, Nb, and Hf using inductively coupled plasma mass spectrometry at Memorial University of Newfoundland. The results are given in Table 1. The method, as well as the quality of the data, was described by Longerich et al. (1990). In general, the precision is 2%–10% for trace elements. Secondary processes that affected the rocks were accompanied by selective chemical modifications. Elevated loss on ignition (LOI) values and variable abundances of K, Rb, and possibly Sr are attributed to alteration and/or metamorphic processes. Some rocks are spilitized, and their abundances of Na2O were
Ordovician–Devonian oceanic basalts in the Cosoltepec Formation
481
TABLE 1. REPRESENTATIVE CHEMICAL ANALYSES OF MAFIC ROCKS IN THE COSOLTEPEC FORMATION AND THREE IGNEOUS ROCKS FROM THE XAYACATLÁN FORMATION Sample (wt%)
COS-3
SiO2 TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5 LOI Total
51.86 2.36 15.01 15.24 0.09 2.53 2.20 6.24 0.24 0.62 2.56 98.94
COS-5 51.60 2.19 15.64 16.07 0.14 2.08 1.70 6.95 0.25 0.88 2.10 99.60
COS-9 49.21 1.08 14.74 10.14 0.15 7.88 9.52 3.88 0.26 0.08 2.53 99.47
COS-16 47.74 1.45 14.49 12.32 0.20 6.67 11.81 2.29 0.22 0.11 2.53 99.84
ppm Cr 358 423 323 88 Ni 126 285 81 73 Co 61 100 46 47 V 270 253 198 255 Cu 44 49 62 97 Pb 5 5 1 5 Zn 90 141 94 110 Rb 11 12 10 7 Ba 5 5 118 49 Sr 73 84 124 165 Ga 15 14 11 18 Ta 2.45 2.34 0.08 0.13 Nb 51.7 49.8 3.4 3.9 Hf 4.34 4.04 1.63 2.22 Zr 210 193 62 83 Y 18 28 20 26 Th 5.08 4.85 0.09 0.15 U 0.00 0.00 0.00 1.00 La 16.90 26.53 1.68 2.67 Ce 43.44 62.54 5.47 7.79 Nd 22.06 25.85 5.61 7.90 Sm 4.64 5.41 2.09 2.76 Eu 1.42 1.85 0.81 1.05 Gd 4.58 5.93 3.08 4.11 Tb 0.69 0.90 0.55 0.75 Dy 4.04 5.33 3.75 5.07 Ho 0.69 0.94 0.73 1.01 Er 1.88 2.51 2.14 2.95 Tm 0.25 0.33 0.32 0.43 Yb 1.49 1.99 2.07 2.78 Lu 0.22 0.29 0.30 0.41 Notes: LOI—loss on ignition; n.d.—not determined.
COS-31
COS-33
COS-44
COS-46
COS-52
IXCA2
46.95 1.31 15.12 11.79 0.20 7.22 10.92 2.75 0.69 0.10 2.95 100.00
47.88 1.33 15.62 9.58 0.26 7.60 9.47 3.51 0.62 0.12 3.89 99.89
47.44 1.15 15.52 10.33 0.19 8.54 9.49 2.08 1.67 0.11 3.33 99.85
47.08 1.82 14.09 13.41 0.15 7.33 9.69 3.08 0.56 0.15 2.68 100.04
47.39 1.23 14.88 10.99 0.19 7.95 11.96 2.53 0.22 0.10 3.44 100.87
46.75 1.43 15.12 11.36 0.15 6.16 10.75 3.77 0.05 0.12 4.51 100.17
306 110 52 238 120 0 65 20 117 196 17 0.23 6.3 2.02 72 18 0.32 0.50 3.83 10.13 7.92 2.41 0.91 3.18 0.56 3.62 0.72 2.06 0.30 1.94 0.28
315 119 54 243 132 0 81 22 218 93 15 0.30 6.5 2.00 74 19 0.33 2.00 4.05 10.56 8.29 2.47 0.93 3.28 0.58 3.85 0.75 2.19 0.31 2.01 0.29
268 117 43 211 57 0 68 34 608 100 16 0.41 2.7 2.18 81 21 0.73 0.50 4.68 11.81 8.88 2.60 0.92 3.49 0.61 4.02 0.82 2.41 0.35 2.29 0.34
154 102 63 299 163 3 90 16 62 215 18 0.40 8.7 2.85 108 22 0.53 1.00 5.97 15.80 12.08 3.53 1.33 4.40 0.74 4.73 0.91 2.58 0.37 2.35 0.33
451 130 49 222 107 3 47 8 133 124 15 0.26 4.7 1.73 64 16 0.29 0.50 3.48 9.04 7.13 2.15 0.83 2.88 0.51 3.29 0.66 1.92 0.27 1.81 0.26
242 108 46 244 103 5 60 5 66 217 19 n.d. 2 n.d. 85 28 <2 <1 12 n.d. 15 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
IXCA5
IXCA3
47.73 1.65 15.23 12.51 0.20 8.03 6.46 3.99 0.40 0.08 3.65 99.93
76.55 0.48 11.05 3.21 0.02 0.79 0.67 1.31 1.97 0.10 2.96 99.11
255 143 67 278 84 3 96 13 903 96 18 n.d. 1 n.d. 93 33 4 1 21 n.d. 23 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
38 8 8 87 5 10 52 109 266 42 13 n.d. 14 n.d. 424 40 14 4 24 n.d. 15 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
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probably changed as well. The most highly altered samples (those with LOI > 7 wt% or containing abundant carbonate veinlets) were omitted from further consideration. To minimize possible alteration effects on rock composition, the evaluation of the petrogenesis and tectonic setting of the mafic rocks is based mainly on trace elements, e.g., high-field-strength elements (HFSE) and the REE considered relatively “immobile” in hydrothermal fluids (Winchester and Floyd, 1977). Whole-Rock Geochemistry The metavolcanic rocks of the Cosoltepec Formation (and two samples of the Xayacatlán Formation) include a variety of mafic
65 Rhyodacite/Dacite SiO2 (wt%)
Trachyte
60
Andesite
Trachyte/ Andesite
55
Phonolite Subalkaline Basalt
50
45
Group 1 Group 2 Ixcamilpa
40 0.001
Basanite/ Trachyte/ Nephelinite
Alkali Basalt
Zr / TiO2
0.01
.1
Figure 3. SiO2 (wt%) versus Zr/TiO2 diagram (based on Winchester and Floyd, 1977) for the mafic rocks of the Cosoltepec Formation.
rocks that can be subdivided according to their composition and location into two distinct groups. The first group, which includes most of the collected samples (COS 9, 10, 15, 16, 17, 19, 24–33, 35, 36, 39, 40, 42–44, and 46–54, as well as IXCA2 and IXCA5), is composed of basalts, whereas the second group (COS 3–8) is made up of basaltic to andesitic rocks. The relationship between the two units has not been observed; they probably belong to two geochemical groups that correspond to different tectonic settings and represent different tectonic slices. These two rock groups are discriminated on the SiO2 versus Zr/TiO2 graph (Fig. 3), the Zr/TiO2 versus Nb/Y graph (Fig. 4), and the Mg# (100 × MgO/ MgO+FeO* in mole %) versus TiO2 diagram (Fig. 5A). The basalts of Group 1 have 46%–52% SiO2 (LOI-free; Fig. 3), and their Mg# values vary between 63 and 35. They display tholeiitic Fe, Ti, and V enrichment trends with increasing differentiation (Fig. 5A, B, and D). Additionally, the rocks show an increase of P and Zr but a decrease of Mg, Cr, and Al with differentiation (Fig. 5C), whereas the Al/Ca ratio remains relatively constant, indicating that the Group 1 rocks were affected by fractional crystallization dominated by plagioclase and clinopyroxene. Compared to primitive island arc tholeiites, the basalts have higher abundances of Ti (TiO2 content is typically >1 wt%; Fig. 5A) and other HFSE and higher Ti/V ratios (32–42; see Shervais, 1982; Fig. 6). The relationships between Ti-Cr-Ni in the Group 1 rocks are also characteristic of nonorogenic basalts (Fig. 7). The REE patterns of the Group 1 volcanic rocks are either flat or depleted in light REE, with (La/Yb)n ~ 0.5–1.5 and (La/Sm)n ~ 0.4–1 (Fig. 8). Their mantle-normalized trace element patterns are relatively flat and do not show a depletion of Nb relative to light REE and Th (Fig. 9). These tholeiitic basalts of the Cosoltepec Formation compositionally resemble MORB derived from heterogeneous sources (Schilling et al., 1983). In contrast, Group 2 (COS-3–8) volcanic rocks are basalts to andesites, with SiO2 ranging between 46 and 61 wt% and Mg# < 25. Unlike typical calc-alkaline volcanic lavas, they have high contents of TiO2 (1.75–2.75 wt%; Fig. 5A) and some other HFSE including Nb (30–40 ppm) and have relatively high contents of transition elements such as Cr (usually 350–450 ppm) and Ni
Andesite 0.01
Zr / TiO2
Andesite/Basalt Alkaline Basalt
SubAlkaline Basalt Group 1 Group 2 Ixcamilpa
Hermanns, 1994
Nb/Y 0.001 0.01
0.1
1
Figure 4. Zr/TiO2 versus Nb/Y diagram (based on Winchester and Floyd, 1977) for the mafic rocks of the Cosoltepec Formation. Also shown is the field of amphibolitic rocks recorded by Hermanns (1994).
0
10
20
Group 1 Group 2 Ixcamilpa
30 Mg #
40
Calcalkaline trend
c
iti
ei
ol
Th
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d
en
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14
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18
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10
20
30 Mg #
40
Calcalkaline trend
Th ol ei iti c
50
Tr en d
Tr en d
Calcalkaline trend
Th ol ei iti c
60
D
B
Figure 5. Diagrams of Mg# (100 × MgO/MgO+FeO* in mole %) versus (A) TiO2 (tot) (wt%), (B) Fe2O3 (tot) (wt%), (C) Cr (ppm), and (D) V (ppm) for the mafic rocks of the Cosoltepec Formation. The lines showing calcalkaline and tholeiitic trends are after Miyashiro (1974).
100
200
300
400
500
1
2
TiO2 (wt %)
Cr (ppm)
20 Fe 2O3 (wt %) V (ppm)
3
A
Ordovician–Devonian oceanic basalts in the Cosoltepec Formation 483
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V (ppm)
500 Group 1 Group 2 Ixcamilpa
10 50
Ti/1000 (ppm)
400
ARC <- 2 0 -> OFB
10
300
200
100
Ocean Floor Basalts Hermanns, 1994
Low Potassium Tholeiites Group 1 Group 2 Ixcamilpa
100 Ti / 1000 (ppm)
Cr (ppm)
0 0
5
10
15
20
Figure 6. V (ppm) versus Ti (ppm) diagram of Shervais (1982) for the mafic rocks of the Cosoltepec Formation. OFB—ocean floor basalts.
N-MORB
100
1000
Figure 7. Ti (ppm) versus Cr (ppm) diagram of Pearce (1975) for the mafic rocks of the Cosoltepec Formation. Also shown is the field of amphibolitic rocks recorded by Hermanns (1994).
10
A
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COS-16
10
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N-MORB
5
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5 11 COS-44 10
E-MO
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RB
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COS-52
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ea
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nd
Ba sa
lts
10 5
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Figure 8. Chondrite-normalized rare earth element abundances in the mafic rocks of the Cosoltepec Formation: (A and B) Group 1 together with the N-MORB (normal mid-ocean ridge basalts) and E-MORB (enriched mid-ocean ridge basalts) of Sun and McDonough (1989) for comparison; (C) Group 2 together with the ocean island basalts of Sun and McDonough (1989) for comparison. Normalizing values after Sun and McDonough (1989). COS—site prefix.
E-M
OR
10 Rock/Primitive Mantle
100
CO
A B
1 B
COS-33
5
COS-32
100
C Oc
ea
nI
CSO-5
10
sla
nd
Ba
sa
lts
COS-3
Th
Nb
La
Ce
Nd
Zr
Hf
Eu Gd Dy Ho Tm Lu Sm Ti Tb Y Er Yb
Figure 9. Primitive mantle-normalized patterns for the mafic rocks of the Cosoltepec Formation: (A and B) Group 1 together with the N-MORB (normal mid-ocean ridge basalts) and E-MORB (enriched mid-ocean ridge basalts) of Sun and McDonough (1989) for comparison and (C) Group 2 together with the ocean island basalts of Sun and McDonough (1989) for comparison. Normalizing values after Sun and McDonough (1989). COS—site prefix.
Ordovician–Devonian oceanic basalts in the Cosoltepec Formation
uia xa q Oa
OCEAN ISLAND MORB OCEANIC LITHOSPHERE
Caltepec Fault
xa q
uia
ACCRETIONARY PRISM
Oa
RH OC EIC EA N
M
S LA OU U T AR RE HER GI NT N N IA
Future Subduction CONTINENTAL LITHOSPHERE Zone
DEVONO-CARBONIFEROUS ALTERNATIVE ORIGINS Derived from upper plate by subduction erosion IXCA MORB PIAXTLA ECLOGITIC SUITE
Derived from lower plate
Figure 10. Cross-sections showing the tectonic interpretations inferred for (A) the deposition of the Cosoltepec Formation and its ocean basalt lenses in the Silurian and (B) the development of an accretionary prism in the Devonian–Carboniferous, at which time tectonic juxtaposition of the Cosoltepec Formation, the ocean lenses, and eclogitic rocks took place. MORB—mid-ocean ridge basalts.
SIBERIA Cadomia
BALTICA
N EA
C
O
NORTH AMERICA lin a
The geochemical affinity of all the mafic rocks within the Cosoltepec Formation indicates extrusion within an ocean basin from either a mid-oceanic ridge (Group 1) or a plume (Group 2). Tectonic interleaving of these oceanic rocks with the Cosoltepec Formation and the Piaxtla Suite is inferred to have taken place during the first stage of deformation immediately prior to deposition of the uppermost Devonian Otates Formation (Vachard and Flores de Dios, 2002; Derycke-Khatir et al., 2005); however, the deformation and metamorphism appear to have continued into the Mississippian (Middleton et al., this volume). Subsequent Permo-Triassic deformation superposed dextral north-south shear zones and south-vergent thrusts on the rocks, obscuring the Devono-Carboniferous structures. However, in zones least affected by the Permo-Triassic deformation, tectonic interleaving of high-grade eclogitic-blueschist rocks of the Piaxtla Suite with rocks of the greenschist-facies Cosoltepec Formation suggests that the Devono-Carboniferous deformation was associated with subduction and exhumation. It is inferred that the Piaxtla Suite was derived from the subducting plate, whereas the ocean floor mafic slices in the Cosoltepec Formation lay on the overriding plate (Fig. 10). The continental nature of some of the rocks in the Piaxtla Suite (Middleton et al., this volume) suggests that the leading edge of the continental margin was being subducted beneath the ocean basin. In this context, the high-grade oceanic amphibolites (IXCA 2 and 5 and those from Olinalá; Hermanns, 1994) may have been derived either from the oceanic lithosphere at the leading edge of the subducting slab or from the overriding slab by subduction erosion that transported the rocks to deeper levels followed by obduction (Fig. 10). The presence of an Ordovician–Upper Devonian ocean basin that was first deformed in the Devonian–Carboniferous is more consistent with the life of the Rheic Ocean than with the Cambrian–Ordovician life span of the Iapetus Ocean. The detrital zircon age suite recovered from the Cosoltepec Formation may be most directly derived from the adjacent Oaxacan Complex and Amazonia (Keppie et al., 2006). Thus, the Cosoltepec Formation and the mafic lava slices are placed outboard of Oaxaquia and Amazonia in a Silurian reconstruction (Fig. 11). Oblique dextral obduction of part of the Rheic Ocean (the Cosoltepec Formation and oceanic basalts) during the Devonian–Carboniferous
RHEIC OCEAN
ro
DISCUSSION AND CONCLUSIONS
COSOLTEPEC CONTINENTAL RISE PRISM
SILURIAN
Ca
(~100–300 ppm). Relative to Group 1 rocks, they have an elevated abundance of incompatible trace elements, higher ratios of Ti/V (46–54 versus the 32–42 of Group 1; Fig. 6), Nb/Y (1.3–1.8 versus 0.02–0.36; Fig. 4), and Zr/Y (6.2–8.6 versus 2.2–5.2). Their REE patterns are distinctly fractionated, with (La/Yb)n ~ 7–8 and (La/Sm)n ~ 2–2.7 (Fig. 8), whereas the mantle-normalized trace element patterns are fractionated but do not show Nb or Ti negative anomalies. Overall, these rocks resemble suites from oceanic islands (Fig. 9). Their chemistry and petrography are similar to those of amphibolitic rocks adjacent to Olinalá (Figs. 1, 4, and 7; Hermanns, 1994).
485
a Av
ni lo
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Figure 12. Devono-Carboniferous paleogeographic map (modified after Keppie, 2004) showing subduction of the leading edge of Oaxaquia/Amazonia beneath the Rheic Ocean. MORB—mid-ocean ridge basalts; OIB—ocean island basalts. X marks the South Pole.
is inferred to have taken place during amalgamation of Pangea (Fig. 12). The presence of Mississippian, midcontinent (USA) brachiopod fauna in Oaxaquia (Navarro-Santillán et al., 2002) suggests that North and Middle America were in close proximity by this time. ACKNOWLEDGMENTS We are grateful to Drs. W. Kramer and F. Ortega-Gutiérrez for their constructive reviews of this article. We would like to acknowledge PAPIIT (Programa de Apoyo a Proyetos de Investigación e Innovación Tecnológica) grant IN103003 to JDK and an NSERC (National Science and Engineering Research Council) Discovery grant to JD that facilitated field work and chemical analysis. We would also like to thank Miguel Morales for assistance with drawing the figures. REFERENCES CITED Derycke-Khatir, C., Vachard, D., Dégardin, J.-M., Flores de Dios, A., Buitrón, B., and Hansen, M., 2005, Late Pennsylvanian and Early Permian chondichthyan microremains from San Salvador Patlanoaya (Puebla, Mexico): Geobios, v. 38, p. 43–55, doi: 10.1016/j.geobios.2003.06.008. Dostal, J., Baragar, W.R.A., and Dupuy, C., 1986, Petrogenesis of the Natkusiak continental basalts, Victoria Island, N.W.T: Canadian Journal of Earth Sciences, v. 23, p. 622–632. Dostal, J., Dupuy, C., and Caby, R., 1994, Geochemistry of the Neoproterozoic Tilemsi belt of Iforas (Mali, Sahara): A crustal section of an oceanic
island arc: Precambrian Research, v. 65, p. 55–69, doi: 10.1016/03019268(94)90099-X. Dostal, J., Keppie, J.D., Nance, R.D., Miller, B.V., and Cooper, P., 2003, Xayacatlán Formation, Acatlán Complex, southern Mexico: Tectonic implications: Libro de Resúmenes, IV Reunion Nacional de Ciencias de la Tierra, Universidad Nacional Autónoma de México, Juriquilla, p. 151. Elías-Herrera, M., and Ortega-Gutiérrez, F., 2002, Caltepec fault zone: An Early Permian dextral transpressional boundary between the Proterozoic Oaxacan and Paleozoic Acatlán complexes, southern Mexico, and regional implications: Tectonics, v. 21, no. 3, doi: 10.1029/200TC001278, doi: 10.1029/2000TC001278. Hermanns, R.L., 1994, Der paläozoiche Acatlán-Komplex bei Olinala, Mexiko: Metamorphose-facies, Geochemie von Metabasalten und Deformationgeschichte [diploma thesis]: Germany, Universität Tübingen, 103 p. Keppie, J.D., 2004, Terranes of Mexico revisited: A 1.3 billion year odyssey: International Geology Review, v. 46, p. 765–794. Keppie, J.D., and Ramos, V.S., 1999, Odyssey of terranes in the Iapetus and Rheic Oceans during the Paleozoic, in Ramos, V.S., and Keppie, J.D. eds. Laurentia-Gondwana connections before Pangea: Boulder, Colorado, Geological Society of America Special Paper 336, p. 267–276. Keppie, J.D., Sandberg, C.A., Miller, B.V., Sánchez-Zavala, J.L., Nance, R.D., and Poole, F.G., 2004a, Implications of latest Pennsylvanian to Middle Permian paleontological and U-Pb SHRIMP data from the Tecomate Formation to re-dating tectonothermal events in the Acatlán Complex, southern Mexico: International Geology Review, v. 46, p. 745–754. Keppie, J.D., Nance, R.D., Powell, J.T., Mumma, S.A., Dostal, J., Fox, D., Muise, J., Ortega-Rivera, A., Miller, B.V., and Lee, J.W.K., 2004b, MidJurassic tectonothermal event superposed on a Paleozoic geological record in the Acatlán Complex of southern Mexico: Hotspot activity during the breakup of Pangea: Gondwana Research, v. 7, p. 238–260, doi: 10.1016/S1342-937X(05)70323-3. Keppie, J.D., Miller, B.V., Nance, R.D., Murphy, J.B., and Dostal, J., 2004c, New U-Pb zircon dates from the Acatlán Complex, Mexico: Implications for the ages of tectonostratigraphic units and orogenic events: Geological Society of America, Abstracts with Programs, v. 36, no. 2, p. 104. Keppie, J.D., Nance, R.D., Fernández-Suarez, J., Storey, C.D., Jeffries, T.E., and Murphy, J.B., 2006, Detrital zircon data from the eastern Mixteca terrane, southern Mexico: Evidence for an Ordovician–Mississippian continental rise and a Permo-Triassic clastic wedge adjacent to Oaxaquia: International Geology Review, v. 48, p. 97–111. Landing, E., Keppie, J.D., and Westrop, S.R., 2006, Lower Paleozoic of northwest Gondwana—terminal Cambrian–lowest Ordovician Tiñu Formation, Oaxaca, southern Mexico: Geological Society of America, Abstracts with Programs, v. 38, no. 2, p. 21. Longerich, H.P., Jenner, G.A., Fryer, B.J., and Jackson, S.E., 1990, Inductively coupled plasma–mass spectrometric analysis of geological samples: A critical evaluation based on case studies: Chemical Geology, v. 83, p. 105–118, doi: 10.1016/0009-2541(90)90143-U. Malone, J.W., Nance, R.D., Keppie, J.D., and Dostal, J., 2002, Deformational history of part of the Acatlán Complex: Late Ordovician–Early Silurian and Early Permian orogenesis in southern Mexico: Journal of South American Earth Sciences, v. 15, p. 511–524, doi: 10.1016/S08959811(02)00080-9. Middleton, M., Keppie, J.D., Murphy, J.B., Miller, B.V., Nance, R.D., OrtegaRivera, A., and Lee, J.K.W., 2007, P-T-t constraints on exhumation following subduction in the Rheic Ocean from eclogitic rocks in the Acatlán Complex of southern México, in Linnemann, U., et al., eds., The evolution of the Rheic Ocean: From Avalonian-Cadomian active margin to Alleghenian-Variscan collision: Geological Society of America Special Paper 423, doi: 10.1130/2007.2423(24). Miller, B.V., Dostal, J., Keppie, J.D., Nance, R.D., Ortega-Rivera, A., and Lee, J.K.W., 2007 (this volume), 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, in Linnemann, U., et al., eds., The evolution of the Rheic Ocean: From Avalonian-Cadomian active margin to AlleghenianVariscan collision: Geological Society of America Special Paper 423, doi: 10.1130/2007.2423(24). Miyashiro, A., 1974, Volcanic rock series in island arcs and active continental margins: American Journal of Science, v. 274, p. 321–355. Navarro-Santillán, D., Sour-Tovar, F., and Centeno-García, E., 2002, Lower Mississippian (Osagean) barchiopods from the Santiago Formation,
Ordovician–Devonian oceanic basalts in the Cosoltepec Formation Oaxaca, México: Stratigraphic and tectonic implications: Journal of South American Earth Sciences, v. 15, p. 327–336, doi: 10.1016/S08959811(02)00047-0. Ortega-Gutiérrez, F., Elías-Herrera, M., Reyes-Salas, M., Macias-Romo, C., and López, R., 1999, Late Ordovician–Early Silurian continental collision orogeny in southern Mexico and its bearing on Gondwana-Laurentia connections: Geology, v. 27, p. 719–722, doi: 10.1130/00917613(1999)027<0719:LOESCC>2.3.CO;2. Pearce, J.A., 1975, Basalt geochemistry used to investigate past tectonic settings on Cyprus: Tectonophysics, v. 25, p. 41–67, doi: 10.1016/00401951(75)90010-4. Ramírez-Espinosa, J., 2001, Tectono-magmatic evolution of the Paleozoic Acatlán Complex in southern Mexico, and its correlation with the Appalachian system [Ph.D. thesis]: Tucson, University of Arizona, 170 p. Sánchez Zavala, J.L., Ortega Gutiérrez, F., and Elías Herrera, M., 2000, La orogenia Mixteca del Devonico del Complejo Acatlán, sur de México: GEOS, Union Geofisica Mexicana, 2nd Reunion Nacional de Ciencias de la Tierra, Resumenes y Programa, v. 20, no. 3, p. 321–322. Schilling, J.G., Zajac, M., Evans, R., Johnston, T., White, W., Devine, J.D., and Kingsley, R., 1983, Petrologic and geochemical variations along the MidAtlantic ridge from 29-degrees-N to 73-degrees-N: American Journal of Science, v. 283, p. 510–586.
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Shervais, J.W., 1982, Ti-V plots and the petrogenesis of modern and ophiolitic lavas: Earth and Planetary Science Letters, v. 59, p. 101–118, doi: 10.1016/0012-821X(82)90120-0. Sun, S.S., and McDonough, W.F., 1989, Chemical and isotopic systematics of oceanic basalts: Implications for mantle composition and processes, in Saunders, A.D., and Norry, M.J., eds., Magmatism in the ocean basins: Geological Society of London Special Publication 42, p. 313–345. Talavera-Mendoza, O., Ruíz, J., Gehrels, G.E., Meza-Figueroa, D.M., VegaGranillo, R., and Campa-Uranga, M.F., 2005, U-Pb geochronology of the Acatlán Complex and implications for the Paleozoic paleogeography and tectonic evolution of southern Mexico: Earth and Planetary Science Letters, v. 235, p. 682–699, doi: 10.1016/j.epsl.2005.04.013. Vachard, D., and Flores de Dios, A., 2002, Discovery of latest Devonian / earliest Mississippian microfossils in San Salvador Patlanoaya (Puebla, Mexico): Biogeographic and geodynamic consequences: Compte Rendu Geoscience, v. 334, p. 1095–1101, doi: 10.1016/S1631-0713(02)01851-5. Winchester, J.A., and Floyd, P.A., 1977, Geochemical discrimination of different magma series and their differentiation products using immobile elements: Chemical Geology, v. 20, p. 325–343, doi: 10.1016/00092541(77)90057-2. MANUSCRIPT ACCEPTED BY THE SOCIETY 3 OCTOBER 2006
Printed in the USA
Geological Society of America Special Paper 423 2007
P-T-t constraints on exhumation following subduction in the Rheic Ocean from eclogitic rocks in the Acatlán Complex of southern México Matt Middleton Department of Earth Sciences, St. Francis Xavier University, Antigonish, Nova Scotia, Canada B2G 2W5 J. Duncan Keppie* Departamento de Geología Regional, Instituto de Geología, Universidad Nacional Autónoma de México, 04510 México D.F. J. Brendan Murphy Department of Earth Sciences, St. Francis Xavier University, Antigonish, Nova Scotia, Canada B2G 2W5 Brent V. Miller Department of Geology, Texas A&M, College Station, Texas 77843 USA R. Damian Nance Department of Geological Sciences, Ohio University, Athens, Ohio 45701 USA Amabel Ortega-Rivera Estación Regional del Noroeste, Instituto de Geología, Universidad Nacional Autónoma, Hermosillo, Sonora, México James K.W. Lee Department of Geological Sciences and Geological Engineering, Queens University, Kingston, Ontario, Canada K7L 3N6
ABSTRACT The Piaxtla Suite of the Acatlán Complex (southern México) has previously been considered a vestige of the Iapetus Ocean that underwent eclogite-facies metamorphism during Late Ordovician subduction and exhumation. Study of granitoid, mafic, and metasedimentary rocks of the Asis Lithodeme of the Piaxtla Suite reveals a complex tectonothermal history involving: (1) eclogite-facies syntectonic metamorphism preserved as aligned omphacite in mafic lenses dated at 346 ± 3 Ma (concordant U-Pb zircon age), which is inferred to result from subduction; (2) polyphase deformation involving WSWENE tectonic transport under amphibolite-facies conditions accompanied by migmatization due to decompression melting dated at ca. 347–330 Ma (SHRIMP [sensitive highresolution ion microprobe] zircon ages); (3) continued deformation under greenschist facies; and (4) development of several phases of late folds and crenulation cleavage. *E-mail:
[email protected]. Middleton, M., Keppie, J.D., Murphy, J.B., Miller, B.V., Nance, R.D., Ortega-Rivera, A., and Lee, J.K.W., 2007, P-T-t constraints on exhumation following subduction in the Rheic Ocean from eclogitic rocks in the Acatlán Complex of southern México, 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. 489–509, doi: 10.1130/2007.2423(25). For permission to copy, contact
[email protected]. ©2007 Geological Society of America. All rights reserved.
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Middleton et al. Pressure, temperature, and time (P-T-t) data suggest rapid isothermal decompression from eclogite to upper amphibolite facies during the Visean (Middle Mississippian) followed by cooling. In the absence of age data for the latter stage, nearby unmetamorphosed latest Upper Devonian sedimentary rocks contain metamorphic, Piaxtla Suite clasts suggesting either diachronism in the exhumation process or synchronous exhumation and subsidence in adjacent areas. Such rapid exhumation of eclogites is typical of continent-continent collision zones, and the subhorizontal, WSW-ENE kinematics is compatible with either lateral thrust ramping or extension in the orogen. Devonian–Carboniferous subduction and exhumation are incompatible with an origin within the Iapetus Ocean, because that ocean had closed by Silurian times. However, they are consistent with oblique subduction of the leading edge of Gondwana along the southern flank of the Rheic Ocean during the amalgamation of Pangea. Keywords: Acatlán Complex, Mexico, eclogite, Carboniferous, Rheic Ocean
INTRODUCTION The Paleozoic Acatlán Complex of southern México contains eclogitic amphibolites that have been inferred to be vestiges of subduction in the Cambro-Ordovician Iapetus oceanic lithosphere (Ortega-Gutiérrez et al., 1999). This proposition was based on the assumption that the ca. 440 Ma megacrystic granites (Esperanza Granitoids) associated with the eclogites represented the products of dehydration melting during exhumation following subduction (Ortega-Gutiérrez et al., 1999; this granite has recently yielded a concordant U-Pb age of 471 ± 6 Ma; Sánchez-Zavala et al., 2004). Conversely, Keppie and Ramos (1999) suggested that the Acatlán Complex formed on the southern margin of the Ordovician–Carboniferous Rheic Ocean. Resolution of this problem is fundamental to Paleozoic continental reconstructions and may be addressed by a detailed field study of these eclogitic amphibolites and associated rocks combined with precise geochronology. This article presents detailed structural, metamorphic, and geochronological data for one of these eclogitic associations, here named the Asis Lithodeme (after its type locality at San Francisco de Asis; Fig. 1). These data record Devono-Carboniferous exhumation following subduction that is more consistent with genesis in the Rheic Ocean than in the Iapetus Ocean. GEOLOGICAL SETTING The Acatlán Complex of the Mixteca terrane is juxtaposed on its eastern side against the ca. 1 Ga Oaxaquia terrane, where the boundary is a Permian dextral flower structure, the Caltepec fault zone (Fig. 1; Elías-Herrera and Ortega-Gutiérrez, 2002). Ortega-Gutiérrez et al. (1999) divided the Acatlán Complex into two major thrust sequences, a lower Petlalcingo Group (the Magdalena Migmatite and the Chazumba and Cosoltepec Formations) and an upper Piaxtla Group, both of which were assumed to be of early Paleozoic age (Fig. 2). These units are unconformably overlain by the Tecomate Formation, which, based on poorly preserved fossils, was assumed to be of Devonian age (Fig. 2;
Ortega-Gutiérrez et al., 1999). However, recent paleontological and geochronological data have shown that the Tecomate Formation is Pennsylvanian–Middle Permian in age, whereas the Magdalena and Chazumba Units (redefined by Keppie et al., 2006, as lithodemes of the Petlalcingo Suite) are of Permo-Triassic age (Fig. 2; Keppie et al., 2006). Correlative sedimentary rocks of the unmetamorphosed Otates and Patlanoaya Formations crop out to the north of the study area and range in age from uppermost Late Devonian to Middle Permian (Vachard et al., 2000; Vachard and Flores de Dios, 2002; Derycke-Khatir et al., 2005). The Magdalena and Chazumba Lithodemes consist of a thick, polydeformed sequence of metapsammites and metapelites (+ calcsilicates and marbles in the Magdalena Lithodeme) that were metamorphosed in the amphibolite facies and migmatized during the Jurassic. The Chazumba Lithodeme additionally contains several tectonic lenses of Jurassic, mafic-ultramafic rocks (Keppie et al., 2004b). The two lithodemes are inferred to be part of a PermoCarboniferous clastic wedge (Keppie et al., 2006). The Cosoltepec Formation structurally overlies the Chazumba Formation and comprises extensive phyllites and quartzites and minor mafic metavolcanic units (Figs. 1B and 2). These rocks have been penetratively deformed three times in the greenschist facies (Malone et al., 2002). The depositional age of part of the Cosoltepec Formation is bracketed between ca. 455 Ma (the youngest concordant detrital zircon; Keppie et al., 2006) and the latest Upper Devonian Otates Formation (Fig. 2; Vachard et al., 2000; Vachard and Flores de Dios, 2002; Derycke-Khatir et al., 2005). The Cosoltepec Formation is structurally juxtaposed against locally eclogitic mafic and ultramafic rocks, high-grade metasedimentary units, granitoid rocks, and migmatites of the Piaxtla Group that are inferred to have been thrust over the Cosoltepec Formation (Ortega-Gutiérrez et al., 1999; Fig. 1B). In view of the inclusion of structurally complex, high-grade metamorphic and metaigneous rocks in the Piaxtla Group, we here propose redefinition of this unit as belonging to the Piaxtla Suite in accordance with the North American Stratigraphic Code (North American Commission on Stratigraphic Nomenclature, 1983). The rocks
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Figure 1. Maps showing (A) the terranes of Middle America and the location of the Mixteca terrane (modified after Keppie, 2004) and (B) the location of the Asis Lithodeme in the Mixteca terrane, made up of the Acatlán Complex (modified after Keppie et al., 2004a).
of the Cosoltepec and Tecomate Formations in sheared contact with the Piaxtla Suite are generally low-grade, suggesting that the Piaxtla Suite had cooled to lower greenschist-facies temperatures before these units were tectonically juxtaposed. Meza-Figueroa et al. (2003) have suggested that nearby eclogitic rocks of the Piaxtla Suite at Piaxtla and Mimilulco represent mid-ocean ridge, ocean island, and island arc basaltic
rocks that were metamorphosed to 11–15 kilobars and 560 ± 60 °C (Fig. 1B), conditions typical of low- to medium-temperature eclogites (Carswell, 1990). Sm-Nd garnet–whole rock data define an isochron of 388 ± 44 Ma. This age is more precisely constrained (386 ± 22 Ma) if strongly retrograded samples are excluded (Yañez et al., 1991). Another garnet–whole rock pair yielded a Sm-Nd isochron age of 416 ± 12 Ma. These ages,
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Figure 2. Time and space diagram for the Piaxtla Suite after OrtegaGutiérrez et al. (1999) and this article. Tr—Triassic; P—Permian; C—Carboniferous; D—Devonian; S—Silurian; O—Ordovician; –C— Cambrian; Fm—Formation.
which span most of the Silurian and Devonian, were interpreted to record postmetamorphic cooling of the rocks through 600 °C (Yañez et al., 1991). Garnet analyses of Meza-Figueroa et al. (2003) plot in the Group C eclogite field of Alpine-type metamorphic rocks and overlap into the field of eclogites in glaucophane schists (Coleman et al., 1965). This Siluro-Devonian metamorphism was inferred to have been followed by partial melting of the metasediments to produce the S-type, megacrystic Esperanza Granitoids (Ortega-Gutiérrez et al., 1999). However, Siluro-Devonian metamorphism is incompatible with recent geochronological results that yielded Ordovician–Early Silurian intrusive ages: (1) a lower intercept U-Pb zircon age of 440 ± 14 Ma and a concordant U-Pb age of 471 ± 6 Ma at the type locality (Ortega-Gutiérrez et al., 1999; Sánchez-Zavala et al., 2004) and (2) a concordant U-Pb zircon age of 478 ± 5 Ma in the western Acatlán Complex (TalaveraMendoza et al., 2005). The sub-greenschist- to greenschist-facies Tecomate Formation consists of conglomerate, sandstone, slate, within-plate mafic and felsic volcanic rocks, and limestone that contains latest Pennsylvanian–Middle Permian conodonts (Keppie et al., 2004a). Zircons from granite pebbles in a conglomerate horizon have yielded SHRIMP (sensitive high-resolution ion microprobe) ages in the range ca. 264–320 Ma. The pebbles were likely derived from the Totoltepec pluton (287 ± 2 Ma, Yañez et al., 1991; 289 ± 1 Ma, Keppie et al., 2004a), which overthrusts the Tecomate and Cosoltepec Formations (Figs. 1B and 2). The Tecomate Formation was penetratively deformed by two sets of structures under greenschist-facies conditions during the PermoTriassic: (1) isoclinal folding associated with north-south dextral
shearing and south-vergent thrusting during which the syntectonic Totoltepec pluton was emplaced and (2) northwest- through north- to northeast-trending upright open folding with an axialplanar crenulation cleavage (Malone et al., 2002; Keppie et al., 2004a). In contrast, the Upper Devonian–Middle Permian conglomerates, sandstones, shales, and limestones of the Otates and Patlanoaya Formations are reported to be only tilted (Vachard et al., 2000). The Tecomate, Otates, and Patlanoaya Formations are reported to rest unconformably on the Piaxtla Suite and the Cosoltepec Formation (Vachard et al., 2000; Sánchez-Zavala et al., 2004). LITHOLOGY, PETROGRAPHY, AND MINERAL CHEMISTRY The study area lies within the northernmost part of the eclogitic Piaxtla Suite between the Piaxtla type area and Mimilulco in the Mixteca terrane (Fig. 1B). The area exposes deformed megacrystic granitoid rocks (lithologically similar to the Esperanza Granitoids) and metasedimentary rocks with amphibolite that are here designated the Asis Lithodeme following the North American Stratigraphic Code (North American Commission on Stratigraphic Nomenclature, 1983). In the east, the Asis Lithodeme and deformed granitoid rocks are in fault contact with rocks of the Tecomate Formation, and elsewhere they are unconformably overlain by Mesozoic and Tertiary rocks (Fig. 1B). The Asis Lithodeme is composed mainly of migmatized metapsammites and metapelites with numerous bands and lenses of amphibolite (see Table 1 for representative microprobe analyses of minerals from the lithodeme). These rocks can be mapped separately from the deformed megacrystic granites, the margins of which are mylonized (Fig. 3). Also present are minor graphitic and chloritic schist and lenses or sheets of amphibolite. Because all contacts are tectonic, it is impossible to ascertain the original relationships between the various rock types. However, the protoliths of the thin bands of amphibolite may have been either minor intrusive sheets or lavas within the metasedimentary rocks. On the other hand, the presence of xenoliths of the metasedimentary rocks within the granitoids suggests an original intrusive relationship. Similarly, the occurrence of quartz-augen granite within the K-feldspar megacrystic granite also suggests an original intrusive relationship, although it is unclear whether these intrusives were part of the same magmatic suite. This quartz-augen granite has yielded a SHRIMP zircon age of ca. 470–420 Ma (Murphy et al., 2006). Metasedimentary Rocks The metasedimentary rocks consist mainly of quartz, plagioclase, biotite, and muscovite (phengite; Fig. 4); minor garnet (mainly almandine with <40% grossular and <20% pyrope; Fig. 4C) and rutile that grow over the main Sa foliation defined by the muscovite and biotite; and secondary chlorite, calcite, and hematite. Migmatization of these metasedimentary rocks has
TABLE 1. (continued) Garnet Phengite Feldspar Sample no.: 87B 87B 87B Analysis no.: 130 131 129 SiO2 37.13 52.84 70.11 0.08 0.48 n.a. TiO2 20.96 28.26 19.62 Al2O3 FeO 15.57 3.25 n.a. MnO 15.19 n.a. n.a. MgO 0.49 2.97 n.a. CaO 11.06 n.a. 0.07 n.a. 0.27 8.65 Na2O n.a. 8.30 n.a. K2O T ot a l 100.49 95.72 98.45 11 oxygens 32 oxygens 24 oxygens Si 5.94 6.88 12.24 Ti 0.01 0.05 n.a. Al 3.95 4.33 4.04 Fe 2.09 0.35 n.a. Mn 2.06 n.a. n.a. Mg 0.12 0.58 n.a. Ca 1.90 n.a. 0.01 Na n.a. 0.07 2.93 K n.a. 1.38 n.a. Total 16.07 13.63 19.21 Notes: 34-11, 34-15, and 42-2—amphibolite lens; 41B—quartz pelitic schist; 42A and 79B—migmatite; 49B—megacrystic granite; 87B—mylonitic granite; n.a.—not available.
TABLE 1. REPRESENTATIVE MICROPROBE ANALYSES (WT%) OF MINERALS FROM THE ASIS LITHODEME, PIAXTLA SUITE, ACATLÁN COMPLEX, SOUTHERN MÉXICO Garnet Amphibole Pyroxene Phengite Feldspar Sample no. 34-11 41B 42A 49B 34-11 42-2 34-11 34-15 42-2 41B 42A 49B 42-2 41B 79B 49B Analysis no. 1320 112 49 208 1332 19 1368 1430 20 110 48 214 22 113 298 208 SiO2 37.98 37.973 37.20 37.55 41.65 42.62 53.43 50.91 48.10 50.06 45.85 51.99 67.53 67.34 68.10 63.55 TiO2 0.23 0.06 0.08 0.08 0.81 1.03 0.34 0.14 0.67 1.36 0.60 0.26 n.a. n.a. n.a. n.a. 20.70 21.12 20.64 21.33 14.13 13.35 0.47 6.87 26.37 26.79 27.98 27.61 19.13 19.91 19.30 17.77 Al2O3 FeO 25.00 28.66 32.49 24.93 16.45 19.92 28.41 13.13 4.28 4.31 4.14 4.90 0.22 n.a. 0.02 n.a. MnO 1.19 1.76 2.88 2.12 0.13 0.24 0.33 0.23 0.01 n.a. 0.02 0.05 n.a. n.a. n.a. n.a. MgO 4.31 3.57 2.97 0.84 9.40 7.67 6.41 14.03 2.74 3.02 2.16 2.65 n.a. n.a. n.a. n.a. CaO 10.95 7.06 4.50 13.78 9.05 8.40 2.15 8.54 0.02 n.a. n.a. n.a. 0.29 0.51 0.18 0.03 n.a. n.a. n.a. 3.46 2.71 6.07 2.89 0.39 0.49 1.16 0.10 11.55 11.36 10.94 0.50 Na2O n.a. n.a. n.a. 0.77 1.06 n.a. n.a. 8.55 9.32 7.57 9.33 0.01 n.a. 0.02 16.48 K2O Total 100.52 100.20 100.74 100.63 95.87 97.01 97.65 96.76 90.48 94.71 88.83 96.25 98.72 99.11 98.56 98.34 24 oxygens 23 oxygens 6 oxygens 11 oxygens 32 oxygens Si 5.95 6.01 5.96 5.95 6.24 6.49 2.10 1.92 6.73 6.713 6.52 6.84 11.97 11.88 12.03 11.99 Ti 0.03 0.01 0.01 0.01 0.09 0.12 0.01 n.a. 0.07 0.137 0.06 0.03 n.a. n.a. n.a. n.a. Al 3.86 3.94 3.90 3.98 2.49 2.40 0.02 0.31 4.34 4.234 4.69 4.28 3.99 4.14 4.02 3.95 Fe 3.27 3.79 4.35 3.30 2.06 2.54 0.93 0.41 0.50 0.483 0.49 0.54 0.03 n.a. n.a. n.a. Mn 0.16 0.24 0.39 0.28 0.02 0.03 0.01 0.01 0.00 n.a. n.a. 0.01 n.a. n.a. n.a. n.a. Mg 1.01 0.84 0.71 0.20 2.10 1.74 0.38 0.79 0.57 0.604 0.46 0.52 n.a. n.a. n.a. n.a. Ca 1.84 1.20 0.77 2.34 1.45 1.37 0.09 0.35 0.00 n.a. n.a. n.a. 0.05 0.10 0.03 0.01 Na n.a. n.a. n.a. n.a. 1.00 0.80 0.46 0.21 0.11 0.127 0.32 0.02 3.97 3.88 3.75 0.18 K n.a.. n.a. n.a. 0.15 0.21 n.a. n.a. 1.53 1.595 1.37 1.57 n.a. n.a. n.a. 3.97 Total 16.10 16.02 16.09 16.05 15.60 15.69 4.00 4.00 13.85 13.894 13.92 13.79 20.02 20.00 19.84 20.11 Continued
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MAP LEGEND Chloritic shear zone
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Figure 3. Geological map of the Asis Lithodeme (Latitude 18°27.6–29.4′; Longitude 98°17.7–19.1′) of the Piaxtla Suite with stereoplots of subareas.
P-T-t constraints on exhumation following subduction in the Rheic Ocean
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Migmatite LEGEND (A-D) Amphibolite Aplite Megacrystic granite Quartz pelitic schist Migmatite (paleasome) Almandine Migmatite (leucosome) Fe3Al2Si3O12 Quartz granite Sodic-calcic amphiboles Mylonitic granite
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Figure 4. Mineral compositions: (A) dioctahedral mica; (B) trioctahedral mica; (C) garnet; (D) plagioclase; (E) amphibole; and (F) pyroxene. Hbd—hornblende.
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produced leucosomes that are predominantly plagioclase (albite; Fig. 4D), quartz, and muscovite (phengite; Fig. 4A). The chlorite and graphite schists appear to be limited to narrow shear zones (Fig. 4). Granitoid Rocks The megacrystic granites are composed of quartz and plagioclase (albite; Fig. 4D), with minor muscovite (phengite) and secondary chlorite. Megacrysts are generally K-feldspar or quartz, with minor garnet (almandine = grossular in the K-feldspar granite, and grossular > almandine in the quartz granite; Fig. 4D). Quartz-augen granite occurs as small bodies within both the K-feldspar granite and the migmatitic metasedimentary rocks (Fig. 4). A few aplitic sheets cut the granites and have essentially the same mineralogy with the addition of secondary epidote. Amphibolites The amphibolites are composed mainly of plagioclase (albite) and aligned amphibole. The amphiboles are mainly sodic-calcic (magnesiotaramite, barroisite, taramite, and ferrobarroisite) or less commonly calcic (ferropargasite and magnesiohornblende; Fig. 4E), with rare omphacite inclusions (Fig. 4F). Also present are epidote, muscovite (phengite; Fig. 4A), biotite (annite/phlogopite; Fig. 4B), and quartz. One or two generations of garnet are composed of almandine > grossular and <20% pyrope (Fig. 4C), with the second generation overgrowing the composite foliation. Hematite, chlorite, titanite, ilmenite, fluorite, and calcite occur as accessory minerals. METAMORPHISM Using mineral chemistry, we have attempted to define the pressure and temperature (P-T) conditions at various stages in the geological history of the Asis Lithodeme. But because the contacts between the various rock units are highly tectonized, we discuss the P-T conditions from each unit separately as it is possible they underwent different P-T paths. Amphibolite Peak metamorphism in the eclogite facies is preserved as rare cores of omphacite-garnet in amphibolite lenses and bands. Garnet compositions fall in the Group C eclogite field of Alpinetype eclogites (Fig. 4A; Coleman et al., 1965). They also overlap the field of eclogites within glaucophane schists (Fig. 4A), suggesting that the P-T path passed through the blueschist field, although no blueschist minerals were observed. Peak temperatures and pressures for the Asis Lithodeme are estimated to be (1) 896–1684 °C using Ellis and Green (1979) and 627–1135 °C using Raheim and Green (1974) for garnet-clinopyroxene pairs (Table 2A) and (2) ~14–30 kb using the molar fraction of jadeite (Table 3). Using coexisting garnet and mica compositions in the
TABLE 2A. GEOTHERMOMETRY OF GARNETCLINOPYROXENE PAIRS FROM THE ASIS LITHODEME, PIAXTLA SUITE, ACATLÁN COMPLEX, SOUTHERN MÉXICO Sample
Garnet-clinopyroxene Gt IN KD T (°C)* Xca 34-11 0.289 0.286 1684 34-15 0.286 1.7616 896 Gt Note: Xca —Molar fraction; KD—equilibrium constant. *Ellis and Green (1979). † Raheim and Green (1974).
†
T (°C) 1135 627
TABLE 2B. GEOTHERMOMETRY OF GARNETPHENGITE PAIRS FROM THE ASIS LITHODEME, PIAXTLA SUITE, ACATLÁN COMPLEX, SOUTHERN MÉXICO Garnet-phengite o Sample no. In KD T ( C)* 42-2-1 1.95 400 42-2-2 1.74 427 42-2-3 1.77 422 42-2-4 1.62 442 42-2-5 1.86 411 42-2-6 1.49 461 42-2-7 1.90 405 42-2-8 1.80 419 42-2-9 2.22 386 42-2-10 4.09 474 101-4-1 2.74 314 101-4-2 2.89 301 101-4-3 2.52 336 101-4-4 2.68 320 101-4-5 2.52 336 101-4-6 2.55 333 101-4-7 2.57 331 101-4-8 2.42 346 101-4-9 2.19 371 Note: KD—equilibrium constant. *Source: Krogh and Raheim (1978).
TWEEQU (Thermobarometry with Estimation of Equilibration State) program of Berman (1991) yields temperatures of ~750– 850 °C (Fig. 5B). Magnesio-taramite and barroisite-hornblende compositions in the plot of Brown (1977) give pressures of ~8 kilobars and ~5.5 kilobars, respectively (Fig. 5C). Retrograde metamorphism is indicated by (1) the compositions of garnetphengite pairs at ~300–474 °C (Table 2B), (2) the Si content per unit of phengite at 3.5–10 kilobars (Table 3), and (3) the magnesio-hornblende compositions at temperatures between ~500 and 350 °C (Fig. 5D). The appearance of secondary epidote and chlorite indicates progressive retrograde metamorphism. Migmatized Metasediments Peak metamorphic conditions in the upper amphibolite facies are indicated by migmatization of garnet-mica schists. Using the phengite compositions in the pelitic rocks indicates pressures of ~7 kilobars at temperatures of ~600–620 °C (Table 4), which is
P-T-t constraints on exhumation following subduction in the Rheic Ocean TABLE 3. GEOBAROMETRY OF MAFIC ROCKS FROM THE ASIS LITHODEME, PIAXTLA SUITE, ACATLÁN COMPLEX, SOUTHERN MÉXICO † Ja/f.u.* Si/f.u. 1 2 Sample no. P (kilobars) P (kilobars) § 34-11 30 n.d. # 34-15 24 n.d. § n.d. 14 42-2-1 n.d. 9.0 42-2-2 n.d. 10.0 42-2-3 n.d. 10.0 42-2-4 n.d. 8.0 42-2-5 n.d. 9.0 42-2-6 n.d. 5.0 42-2-7 n.d. 9.0 42-2-8 n.d. 10.0 42-2-9 n.d. 8.5 42-2-10 n.d. 5.0 101-4-1 n.d. 6.0 101-4-2 n.d. 3.5 101-4-3 n.d. 6.5 101-4-4 n.d. 6.5 101-4-5 n.d. 4.0 101-4-6 n.d. 4.0 101-4-7 n.d. 6.5 101-4-8 n.d. 3.5 101-4-9 n.d. 7.0 Notes: 34-11, 34-15, 42-2, and 101-4—amphibolitic lenses; n.d.—not determined. † † * Molar fraction of Ja (*) and Si ( ) content per formula unit of phengite. § Temperature calculated from garnet-clinopyroxene by Raheim and Green (1974). # Temperature calculated from garnet-clinopyroxene by Ellis and Green (1979). 1 Source: Holland (1980) (pressures are overestimated by about 1.5–2.0 kilobars due to Holland’s (1980) pressure correction for impure jadeite). 2 Source: Massonne and Schreyer (1987).
consistent with the presence of migmatization. Kyanite, sillimanite, and andalusite were not encountered. The presence of chlorite indicates retrograde metamorphism. Metagranitoids Peak metamorphism in the granitoid rocks is difficult to estimate due to the lack of index minerals; however, the presence of almandine garnet and the absence of orthopyroxene indicates metamorphism in the amphibolite facies. Metamorphic conditions of ~5 kilobars at 450–550 °C in the granitoid rocks are estimated using phengite compositions (Table 5), i.e., in the upper greenschist facies. Secondary chlorite again indicates retrograde metamorphism. STRUCTURE The Asis Lithodeme displays a complex structural history that developed under decreasing metamorphic grades from eclogite facies through amphibolite- and greenschist-facies and
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into sub-greenschist conditions. We designate the structures at each grade as De (eclogite facies), Da (amphibolite facies), Dg (greenschist facies), and Dlate (sub–greenschist facies). Where several sets of structures formed in one facies, a further subscript is added, viz., Da1, Da2, and Da3, to denote three sets of structures formed under amphibolite-facies conditions. Eclogite Facies (De) The earliest structure in the Asis Lithodeme is preserved only in the cores of amphibolite lenses where omphacite needles within amphibole are aligned, suggesting that they formed a foliation (Se1) developed at eclogite facies of metamorphism. Amphibolite Facies (Da) Psammiic Rocks Several generations of folds associated with migmatization in the Asis Lithodeme were produced under amphibolite-facies metamorphic conditions. This is best observed in the metasedimentary rocks at locality 1 (Fig. 3), where polyphase deformation of psammitic bands interlayered with thin pelitic bands is preserved (Fig. 6). The earliest folds (Fa1) are tight-isoclinal and have an axial-planar foliation (Sa1) defined by aligned mica. These folds deform a compositional banding inferred to be of tectonic origin (possibly Se). In the limbs of these folds, Se and Sa1 are indistinguishable and form a composite foliation. This composite foliation is deformed by isoclinal folds (Fa2) that are refolded by tight isoclinal folds (Fa3), producing a mushroomtype interference pattern (Fig. 6). The Fa1, Fa2, and Fa3 fold axes are nearly parallel to a strong lineation defined by stretched quartz, which is inferred to represent the finite stretching lineation (La1 + 2 + 3) produced during the deformation that generated the three sets of folds (Fig. 6). Structural data here suggest that the Fa3 fold has top-to-the-northeast kinematics. These Fa2 and F3a folds are enveloped between two discrete shear zones (Fig. 6) on which occurs a strong northeast-plunging stretching lineation that is parallel to the composite stretching lineation in the study area, suggesting that the shear zones were active during the development of the folds. Examination of an Fa2 fold just north of these shear zones indicates that it is a sheath fold with a northeast-plunging fold axis that is parallel to a strongly developed lineation defined by stretched quartz. That all of these folds developed under amphibolite facies during a single phase of progressive deformation is suggested by the observation that although some migmatitic leucosomes are parallel to the Sa1 foliation and are deformed by Fa2 and Fa3, others are parallel to the axial planes of Fa3 folds. Pelitic Rocks The pelitic metasedimentary rocks in the Asis Lithodeme are more tightly folded and more extensively migmatized than other lithologies. Rarely, hinges of isoclinal folds with the Sa1 foliation parallel to the axial planes of F1a were observed. More
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Figure 5. Mineral data from mafic rocks in the Asis Lithodeme, Piaxtla Suite, Acatlán Complex, plotted on various diagrams: (A) ternary diagram for garnet (fields of Coleman et al., 1965); (B) garnet-biotite compositions plotted on the pressure-temperature graph of Berman (1992); (C) amphibole compositions plotted on the diagram of Brown (1977); (D) pargasitic and magnesiohornblende compositions plotted on a plagioclase-amphibole geothermometer after Spear (1980). TABLE 4. GEOTHERMOBAROMETRY OF METASEDIMENTARY ROCKS IN THE ASIS LITHODEME, PIAXTLA SUITE, ACATLÁN COMPLEX, SOUTHERN MÉXICO † † Sample no. Si/f.u.* P (kilobars) T (°C) 41b 3.4 12 600 41b 3.3 9 600 71a 3.3 9 600 71a 3.4 12 610 71a 3.2 7 600 71a 3.3 9.5 610 42a 3.3 9.5 600 42a 3.3 9.5 605 42a 3.3 9.5 610 42a 3.3 9.5 615 42a 3.3 10 620 79a 3.3 9 600 79a 3.3 9.5 605 79a 3.4 11 610 79a 3.4 11.5 615 79a 3.3 10 620 79b 3.4 12 600 79b 3.4 12 605 79b 3.3 9 610 79b 3.3 9.5 615 79b 3.4 12 620 Notes: 41b and 71a—quartz pelitic schist; 42a—migmatite; 79a—migmatite paleosome; 79b—migmatite leucosome. *Si content per formula unit of phengite. † Source: Massonne and Schreyer (1987).
TABLE 5. GEOTHERMOBAROMETRY OF GRANITOID ROCKS IN THE ASIS LITHODEME, PIAXTLA SUITE, ACATLÁN COMPLEX, SOUTHERN MÉXICO † † T (°C) Sample no. Si/f.u.* P (kilobars) 49b 3.4 10.5 490 49b 3.4 11.0 510 49b 3.5 13.0 530 49b 3.5 13.5 550 84-3 3.2 5.0 450 84-3 3.5 12.5 470 84-3 3.5 13.0 490 84-3 3.4 11.0 510 84-3 3.5 13.5 530 84-3 3.5 14.0 550 84-3 3.4 12.0 570 84-3 3.4 12.5 590 87b 3.2 5.5 490 87b 3.5 13.5 510 87b 3.5 14.0 530 87b 3.5 14.0 550 Notes: 49b—megacrystic granite; 84-3—quartz granite; 87b— mylonitic granite. *Si content per formula unit of phengite. † Source: Massonne and Schreyer (1987).
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Fa2 data
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Fa3 data Axial plane (Fa1) Axial plane (Fa2)
Fold axis (Fa2) Composite stretching lineation
Composite foliation (Se/a1+2)
X X
Composite stretching lineation
Axial plane (Fa3) Fold axis (Fa3)
X X
Fold axis (Fa1)
Composite foliation (Se/a1) Foliation (Se?)
Th
rus
t la
NW
ter
al
Fa1
ram
p Fa3
Vertical section (looking NE)
Motion direction Toward Away Fa2
5m
SE
Fa1 Fa2
Fa3 Fa3
Fa3 Fa3
Figure 6. Stereoplots, sketch, and photographs of structural data from locality 1 (for location see Fig. 3) showing polyphase, amphibolite-facies deformation in metapsammite of the Asis Lithodeme.
500
Middleton et al.
commonly, tight to isoclinal coaxial folds (Fa2 and Fa3) are present. The composite foliation (Se/a1) is defined by aligned muscovite, biotite, and albite that contains aligned muscovite inclusions oblique to the external fabric. Garnet and rutile porphyroblasts generally cut across the composite foliation, indicating that they postdate the foliation. Detailed measurements of a tight Fa2 fold at locality 2 in the pelitic metasedimentary rocks (Fig. 7) reveal that the composite foliation is deformed about a fold axis that is parallel to the stretching lineation defined by stretched quartz. At the southern end of this locality, an amphibolite band is boudinaged in association with the development of a C-S fabric and a gently ENE-plunging stretching lineation, which indicate top-tothe-east kinematics during Fa2 (Fig. 7). In most outcrops, migmatitic leucosomes are parallel to the composite foliation (Se/a1) and are folded about isoclinal Fa2 folds; however, thin migmatitic leucosomes were also observed cutting across the composite foliation and hinges of Fa2 and Fa3 folds, indicating that migmatization continued through Fa1, Fa2, and Fa3 (Fig. 8A). Graphitic Schist Highly deformed graphitic schist occurs as isolated lenses in the northwestern part of the map area (Fig. 3), where they occur along the mylonized boundary between migmatized metasedimentary rocks and either mylonized granite or amphibolite. The schist displays isoclinal folds with easterly plunging fold axes that are parallel to Fa2 and Fa3 folds in the neighboring rocks. The graphite-defined foliation is also folded about a moderately ENE-plunging fold axis that is correlated with the Dlate folds (see later). The high ductility of graphite is inferred to have localized the strain. Amphibolite Fabrics in the amphibolite bands and lenses consist mainly of a foliation that is parallel to the composite Se/a1 foliation in the surrounding metasedimentary rocks and a mineral lineation defined by aligned amphibole that is slightly oblique to the stretching lineation in the surrounding metasedimentary rocks. However, within the centers of the amphibolite bands and lenses, an earlier foliation also defined by aligned amphibole is sometimes preserved, and rarely this foliation is deformed by nearly isoclinal folds that can be correlated with Fa2. Garnet porphyroblasts generally overgrow the Se/a1 composite foliation, but on rare occasions an earlier generation of garnet containing randomly oriented inclusions was observed around which the composite foliation wraps. At one locality (Fig. 8B), migmatitic leucosomes cut the Sa1 foliation in an amphibolite lens and are isoclinally folded, with an axial-planar foliation defined by aligned amphibole. GRANITOIDS In the central parts of the granitoid bodies, a strong foliation defined by aligned muscovite and stretched quartz and albite wraps around megacrysts of deformed K-feldspar and rare garnet. These observations indicate deformation at temperatures >500 °C (assuming Pfluid = Ptotal; Yassaghi, 2000) and indicate that the folia-
tion developed under amphibolite-facies conditions. This foliation is associated with a strong lineation defined by stretched quartz. Migmatitic leucosomes are observed parallel to and cutting across the foliation in the quartz granite. Detailed measurements of granitic sheets within the metasedimentary rocks at locality 3 (Fig. 9) reveal that the foliation in the granitoid sheets is deformed by close, asymmetric Fa3 folds whose axes are parallel to a strongly developed, north-plunging stretching lineation. At this locality, megacrystic K-feldspar σ structures generally show top-to-thesouth kinematics except at the northernmost end, where, they indicate top-to-the-north kinematics. The opposing kinematics suggests that the granitoid body acted as a more competent lens within the metasediments and was squeezed out toward the south during the deformation. This explanation may also explain the opposing kinematics observed in the area as a whole. Greenschist-Facies Fabrics (Dg) Toward the margins of the granitoid bodies, the foliation becomes progressively more mylonitic. In thin section, the mylonitic fabric is defined by stretched quartz that wraps around K-feldspar megacrysts and plagioclase, both of which show brittle deformation suggesting temperatures between 300° and 500 °C (assuming Pfluid = Ptotal: Yassaghi, 2000), suggesting deformation under greenschist-facies conditions. Although the implied grade of metamorphism is lower than that associated with the amphibolite-facies deformation, the orientation of the fabric elements appears to be congruent, suggesting that they represent a continuous phase of deformation that took place during gradually decreasing temperatures. A chloritic shear zone observed in a stream section in the south central part of the study area (Fig. 3) suggests shearing in the lower greenschist facies. It is parallel to the foliation in the surrounding rocks and was not encountered elsewhere. Late Folds and Crenulation Cleavage (Dlate) ENE, Steeply Inclined Folds At locality 2 (Figs. 3 and 7), a late, moderately WSW-plunging, steeply inclined cylindrical antiform with an axial-planar crenulation cleavage deforms a megacrystic granite sheet within the migmatized metasediments. This fold deforms the stretching lineation into a small circle pattern, indicating a flexural folding mechanism. At locality 4 (Fig. 10), a late box fold has ENEstriking axial planes that dip both north and south. The fold axes of the box fold are nearly parallel and plunge gently ENE: the intersection between the axial planes is slightly oblique to the composite foliation planes. The conjugate nature of the box axial planes indicates subhorizontal, north-south shortening and vertical extension. Other folds of similar orientation include (1) a late fold in the graphite schist locality (Fig. 3), (2) one of the crenulation cleavages (Fig. 11), and (3) a macroscopic fold located along the eastern margin of the area (Fig. 3). These Dlate folds are approximately parallel to the Fa2 and Fa3 folds and so contribute
P-T-t constraints on exhumation following subduction in the Rheic Ocean Granitic sheet sample 102-3a Composite foliation (Se/a1) Stretching lineation
N Migmatitic gneiss
Fault
25 24
Small circle about the Phi-axis
55
33 16
65
28 73
52
501
42
45 23 54 37 33 69 26 58 60 21
08
63 52
04 46
47 71 73
84
38 06
85
ane
g pl
tin isec
28
Late
N
fold
b
Figure 7. Map and stereoplots of structural data from locality 2 (for location see Fig. 3) showing an amphibolite-facies fold and a late fold with crenulation cleavage in the migmatized metasediments with amphibolite boudins and a megacrystic granitoid sheet in the Asis Lithodeme.
Phi-circle axis (Late fold axis) 66 59
Map legend 66 Composite foliation 02 Stretching lineation Top to the north Synform Antiform Amphibolitic boudin
Migmatitic gneiss
Fa2 fold axis
63
5m 02 81
77 79
Fa2 folds in migmatitic leucosome cutting amphibolite and Sa1
Se/a1
Sa2 axial-planar foliation
Fa2 Sa1 foliation
Leucosome
A
Amphibolite
B
Figure 8. Photographs of (A) migmatitic pelitic gneiss showing a boudinaged leucosome parallel to the composite foliation Sa1/2 and cross-cutting leucosomes that are gently folded with axial planes parallel to the composite foliation, and (B) amphibolite lens showing foliation, Sa1, truncated by a leucosome that is isoclinally folded by the Fa2 fold with an axial-planar foliation, Sa2.
Fa3 data
502
Middleton et al.
A
N
32 30 28
27
24
5m
B 21
26 27 Composite Foliation (Sa1/2) Fold axis (Fa3) Stretching lineation (La) Crenulation axis
17
65 53 25 18 18 17 68 C 29 67 53 34 22 Foliation 10 22 59
24 05
10 53
D
Stretching lineation Crenulation
72 22 16 13 09 41
15
56
E
kinematics 09 Top to the south kinematics
Figure 9. Map, stereoplots, and map of structural data from locality 3 (for location see Fig. 3) showing stretching lineations coaxial to folds in the megacrystic metagranitoid of the Asis Lithodeme.
Fa3
22
53
30 Top to the north
23
Looking northward
XC Composite foliation Late fold X axial plane Late fold axis
C
A
A A C
1m
A
C
X
Figure 10. Photograph and stereoplots of structural data from locality 4 (for location see Fig. 3) showing a box fold in migmatized metasediments of the Asis Lithodeme.
P-T-t constraints on exhumation following subduction in the Rheic Ocean
503
the direction of expulsion of the granitoid sheet would be gently upward to the WSW.
X XX
U-PB GEOCHRONOLOGY Analytical Methods
X Crenulation cleavage axial plane Crenulation cleavage fold axis Late fold axial plane pole X
Figure 11. Stereoplot summarizing orientations of crenulation cleavages.
to the great circle distribution of composite foliation poles apparent in many of the subarea stereoplots (Fig. 3). ENE and ESE Crenulation Cleavage Two other steeply dipping crenulation cleavages occur sporadically throughout the area and strike ENE and ESE (Fig. 11). In the metasedimentary rocks, they deform muscovite and biotite, which are not recrystallized. In the granitoid rocks, they occur only in the mylonites, where they deform feldspar and quartz, the latter showing recrystallization. These observations suggest that the crenulation cleavages formed at low grades of metamorphism in the sub-greenschist facies. The orientations of the crenulation axes lie in crenulation cleavage planes, and their variable orientation is due to the attitude of the pre-existing foliation surfaces. Intersections between the various crenulation cleavages were not observed, so their relative timing remains unknown. NNW Folds A major NNW-striking fold may be observed in the map pattern and in the stereoplot of the composite foliation in its hinge zone (Fig. 3). The fold appears to pass from an antiform in the north to a synform in the south and may be responsible for repetition of the recumbently folded psammitic rocks. Unfolding Late Folds Most of the late folds are gentle and allow one to unfold them, which results in an ENE-striking, moderately N-dipping composite foliation and a gently ENE-plunging stretching lineation. Based on the K-feldspar and sheath fold kinematics,
Two samples were collected for U-Pb isotopic analysis: an eclogitic amphibolite sample (34-Z; UTM 74.767, 39.284) and the leucosome in the migmatitic metasedimentary rocks (CET-4; UTM 73.805, 39.394; Fig. 3, Table 6). The eclogitic amphibolite is composed of amphibole (magnesiotaramite, taramite, barroisite, and ferrobaroisite) with inclusions of omphacite and overgrowths of rutile and garnet (Alm60–75, Gro20–35, Pyo10–20, Spes0–10), plagioclase (albite), epidote, minor chlorite, biotite, muscovite (annite or phlogopite and phengite), and quartz, with accessory hematite, ilmenite, titanite, and fluorite. The leucosome consists of quartz and albite with minor phengitic muscovite. The samples were crushed, the zircons were separated and processed, and sample 34-Z was analyzed in the Radiogenic Isotope Geochemistry Laboratory at the University of North Carolina–Chapel Hill using the techniques and equipment described in Ratajeski et al. (2001). The zircon analyses were conducted on single grains or a single fragment broken from a grain; the smallest zircons required fractions of 2–5 grains (Table 6). Selected zircon grains from both samples were mounted in epoxy, polished, imaged in cathodoluminescence, and photographed under reflected light, then gold-coated prior to analysis. Uranium-lead isotopic analyses were conducted on the SHRIMP-RG (sensitive high-resolution ion microprobe–reverse geometry) ion microprobe at Stanford University following techniques described in DeGraaff-Surpless et al. (2002). Uranium and thorium concentrations were measured relative to the CZ3 zircon standard, and isotopic ratios and concentrations were measured relative to the R33 zircon standard from a quartz diorite of the Braintree Complex, Vermont (J. Wooden, personal commun., 2006). The data were reduced with the SQUID v. 1.02 program of Ludwig (2001) and plotted with Isoplot Ex, v. 3.00 (Ludwig, 2003). The errors in Table 1 are quoted at 1σ, and the ellipses on the concordia diagrams are shown at the 68.3% confidence level. See the notes to Table 6 for additional analytical details. Results TIMS (thermal ionization mass spectrometry) analyses of zircon fractions from the eclogitic amphibolite (sample 34-Z) yielded one concordant analysis with an age of 346 ± 3 Ma (Table 6, Fig. 12A) and several more or less discordant points that plot on a chord from 346 ± 3 Ma to −62 ± 180 Ma (Fig. 12A). SHRIMP analyses overlap the TIMS data at ca. 346 Ma, and also provide slightly discordant data, with 207Pb/206Pb ages of 840 Ma and 1059 Ma (Fig. 12A, Table 6). SHRIMP analyses of zircons from the migmatitic leucosome (sample CET-4) fall into two clusters (Fig. 12B, Table 6):
Wt. (mg)1
Total1 U (ng) Total2 Pb (pg)
Total2 common Pb (pg) 6 5 5 3 2 4 3 2 7 2
Pb2 Pb
2657 1383 1721 513 370 263 1495 159 609 191
204
206
Pb3 Pb
0.00 0.000 0.000 0.000 0.001 0.001 –0.000 –0.002 –0.002 –0.000
208
206
Pb3 U
0.05519 0.05500 0.05497 0.05456 0.05425 0.05379 0.05378 0.05137 0.05033 0.04908
238
206
0.266 0.909 0.218 0.687 0.641 0.815 0.148 1.440 0.355 1.155
0.40613 0.40598 0.40455 0.40574 0.40261 0.39924 0.39652 0.38861 0.37089 0.36129
Atomic ratios 4 207 Pb3 235 U
% error
0.397 1.004 0.344 1.294 1.088 1.384 0.248 3.669 0.822 1.868
% error4
Sample-spot no.
Atomic ratios 206 206 6 207 U Th % error6 Pb* Pb3 % error Pb3 238 235 (ppm) (ppm) U U (ppm) Asis eclogite: SHRIMP (sample 34-Z: UTM 74.767, 39.284) 34Z-1 285.00 49.80 38.0 0.15420 3.3 1.58700 4.8 34Z-2 18.00 0.06 0.8 0.05240 5.3 0.45000 50.0 34Z-3 1.31 0.01 0.1 0.06570 15.0 1.36000 58.0 34Z-4 93.00 0.31 4.5 0.05600 3.3 0.45400 5.4 34Z-5 174.00 0.06 8.9 0.05900 3.5 0.47100 8.3 34Z-6 213.00 74.20 39.7 0.21520 3.2 2.33600 4.0 34Z-7 25 0.05 1.2 0.05500 3.8 0.42800 9.3 34Z-8 3.89 0.01 0.2 0.04480 13.0 34Z-9 8.14 0.02 0.4 0.05850 3.8 0.65500 11.0 34Z-10 97.00 0.04 4.5 0.05432 1.1 0.38800 3.8 34Z-11 2.37 0.01 0.1 0.02600 47.0 34Z-12 21.00 0.14 1.0 0.05490 2.0 0.47200 6.5 34Z-13 60.00 0.09 2.8 0.05423 1.3 0.44200 4.5 34Z-14 72.00 0.09 3.5 0.05675 1.3 0.52500 4.2 34Z-15 3.76 0.02 0.2 0.05150 10.0 Asis Migmatite: SHRIMP (sample CET-4: UTM 73.805, 39.394) CET4-1 376 196.3 78.7 0.24330 0.460 3.15400 0.850 CET4-2 128 1.0 5.8 0.05239 1.000 0.37400 5.800 CET4-3 260 2.6 12.5 0.05562 1.400 0.39500 3.100 CET4-4 211 84.2 47.1 0.25920 0.620 3.40900 1.100 CET4-5 468 171.9 74.8 0.18590 1.000 2.01300 1.200 CET4-6 574 4.1 27.4 0.05563 0.470 0.41990 1.700 CET4-7 152 44.1 27.4 0.20980 1.100 2.30600 1.700 CET4-8 224 45.2 35.8 0.18570 0.560 1.97300 1.200 CET4-9 254 72.3 39.1 0.17830 0.720 1.81000 2.200 CET4-10 153 13.5 20.8 0.15880 0.700 1.58400 1.700 CET4-11 142 0.3 6.6 0.05414 1.400 0.37400 4.400 CET4-12 159 1.9 7.6 0.05528 0.990 0.39700 3.900 1 Weight estimated from measured grain dimensions and assuming density = 4.67g/cm3; ~20% uncertainty affects only U and Pb concentrations. 2 Corrected for fractionation (0.12 ± 0.08%/atomic mass unit [amu], Faraday-Daly; 0.20 ± 0.1%/amu, Daly) and spike. 3 Corrected for fractionation, blank, and initial common Pb. 4 ID-TIMS (isotopic dilution thermal ionization mass spectrometry) errors are 2σ. 5 207 Pb/235U-206Pb/238U correlation coefficient of Ludwig (1989). 6 SHRIMP (sensitive high-resolution ion microprobe) errors are 1σ.
124 95 108 52 38 72 67 51 150 53
U Pb (ppm) (ppm) Pb3 Pb
Pb3 Pb
6.2 4.2 4.0
0.0624 0.0592 0.0671
0.710 5.800 2.700 0.870 0.740 1.700 1.300 1.000 2.100 1.600 4.200 3.800
10.0 3.6
0.0813 0.0518
0.09403 0.05170 0.05150 0.09541 0.07854 0.05475 0.07970 0.07705 0.07360 0.07230 0.05010 0.05210
3.5 49.0 56.0 4.2 7.6 2.4 8.5
% error6
0.292 0.407 0.255 1.066 0.832 1.071 0.191 3.192 0.701 1.390
% error4
0.0746 0.062 0.151 0.0588 0.0579 0.0787 0.0564
06
207
0.05337 0.05353 0.05337 0.05394 0.05383 0.05383 0.05347 0.05487 0.05345 0.05339
206
207
TABLE 6. U-Pb ISOTOPIC DATA FOR THE ASIS LITHODEME, PIAXTLA SUITE, ACATLÁN COMPLEX, SOUTHERN MÉXICO
Asis eclogite: ID-TIMS (sample 34-Z: UTM 74.767, 39.284) 1, large stubby prism (1) 0.007 0.87 43.20 1.18 2, large flat prism frag. (1) 0.004 0.38 19.00 1.00 3, large prism (1) 0.007 0.76 37.60 1.59 4, med. stubby prism (1) 0.003 0.16 7.76 1.13 5, med. stubby prism (1) 0.003 0.11 5.60 1.15 6, small flat prism (1) 0.002 0.14 7.02 2.07 7, large stubby prism (1) 0.008 0.54 26.30 1.28 8, small thin prisms (2) 0.001 0.05 2.37 1.22 9, large thin prism (1) 0.002 0.30 13.70 1.67 10, small equant (2) 0.001 0.05 2.34 0.98
Analysis no., fraction (number of grains) Pb U
346.3 345.2 345.0 342.4 340.5 337.8 337.7 322.9 316.5 308.8
238
206
Pb Pb
207
344.6 351.4 344.6 368.4 363.9 363.8 348.8 407.0 347.8 345.5
206
1404 329.2 348.9 1486 1099 349.0 1228 1098 1058 950.2 339.9 346.8
925.0 329.0 410.0 352.0 370.0 1257 345.0 282.0 366.0 341.0 163.0 344.4 340.4 355.8 324.0
1509 274.0 261.0 1536 1160 402.0 1190 1122 1030 995.0 201.0 290.0
688.0 573.0 840.0
1229 276.0
1059 680.0 2353 558.0 525.0 1165 470.0
Ages (Ma) 207 Pb Pb 206 U Pb 238
206
346.1 346.0 344.9 345.8 343.5 341.1 339.1 333.4 320.3 313.2
Ages (Ma) 207 Pb 235 U
0.54 0.17 0.45 0.58 0.80 0.27 0.64 0.48 0.32 0.41 0.32 0.25
0.31 0.30 0.30
0.34 0.29
0.69 0.11 0.26 0.62 0.42 0.80 0.41
Rho5
0.68 0.91 0.67 0.57 0.65 0.63 0.64 0.51 0.53 0.67
Rho5
504 Middleton et al.
P-T-t constraints on exhumation following subduction in the Rheic Ocean
A
Amphibolitized Eclogite (34-Z)
(1) generally reversely discordant data with 1σ errors that overlap the concordia between ca. 330 Ma and 348 Ma and (2) nearly concordant data with 207Pb/206Pb ages ranging from ca. 995 Ma to ca. 1536 Ma.
350
2
3
(and -62 ± 180 Ma)
4
340
MSWD = 0.77
5
0.053
Interpretation
6
7
330
206
Pb/
238
U
1
346.3 ± 3.2
0.055
U-Pb zircon ages for the eclogite and the migmatitic leucosome are inferred to indicate cooling through ~800 °C (Heaman and Parrish, 1991), which is similar to the peak temperatures of ~750–850 °C estimated for the Asis Lithodeme. Thus, the concordant ages of 346 ± 3 Ma and ca. 330–347 Ma recorded in the eclogitic amphibolite and the migmatitic leucosome are inferred to closely postdate peak metamorphism. On the other hand, the ages of ca. 950–1500 Ma are interpreted to record inheritance either in the magma source region (eclogite) or in the surrounding metasediments.
8
320
0.051
9 310 0.049
10 300 data-point error ellipses are 2sigma
0.047 0.34
0.36
0.38
207
B
505
Pb/
0.40
235
0.42
U
0.085
Amphibolitized Eclogite (34-Z)
40
(SHRIMP Data)
Ar/ 39Ar GEOCHRONOLOGY
480
0.075
Analytical Methods
0.065
400
206
Pb/
238
U
440
360 1300
0.055
320
1100 900
280
0.045
700 500
0.08
300 0.04
data-point error ellipses are 68.3% conf
0.035 0.25
0.35
0.45
207
C
235
Pb/
0.00 0.0
0.4
0.55 0.8
0.65
U
Migmatitic Leucosome (CET-4)
0.059
370
(SHRIMP Data) 360
206
Pb/
238
U
0.057
350 0.055
340 1400
0.053
330 1000
Muscovite mineral grains were separated from a granitic sheet at locality 2 (Figs. 3 and 7). The grains were pretreated and concentrated by standard techniques and later selected by hand-picking under a binocular microscope from fractions that ranged in size from 40 to 60 mesh at the mineral separation laboratory at UNICIT (Unidad de Investigación en Ciencias de la Tierra)–Universidad Nacional Autónoma de México, Campus-Juriquilla, Querétaro. Mineral separates were loaded into aluminum foil packets and irradiated together with sample Hb3 gr (1072 Ma) as a neutronfluence monitor at the McMaster Nuclear Reactor, Hamilton, Ontario. 40Ar/39Ar analyses were performed by the standard laser step-heating techniques described in detail by Clark et al. (1998) at the Geochronology Research Laboratory of Queen’s University, Kingston, Ontario. The data are given in Table 7 and plotted in Figure 13. All data have been corrected for blanks, mass discrimination, and neutron-induced interferences. For the purposes of this article, a plateau age is obtained when the apparent ages of at least three consecutive steps, comprising a minimum of 55% of the 39Ark released, agree within 2σ error with the integrated age of the plateau segment. The errors shown in Table 7 and on the age spectrum represent the analytical precision at ± 2σ.
320
0.051
Results
600 0.08
310 0.049 0.35
data-point error ellipses are 68.3% conf
0.04 200
0.37
0.39
0.41
0.00 0
207
Pb/
235
0.43
1
U
Figure 12. U-Pb isotopic analyses of zircon plotted on concordia diagrams: (A) eclogitic amphibolite (34-Z) and (B) migmatitic leucosome (CET-4). MSWD—mean square of weighted deviates; SHRIMP—sensitive high-resolution ion microprobe.
40
Ar/39Ar analyses of muscovite from the granitic sheet (sample 102–3a) yielded slightly discordant data, with a plateau age of 351 ± 2 Ma (Fig. 13). Interpretation Although the incorporation of excess Ar in eclogite-facies minerals is often problematic in the interpretation of their
506
Middleton et al. 39
40
Power 0.75 1.25 1.50 1.75 2.00
TABLE 7. Ar/ Ar ANALYSES OF MUSCOVITE IN MEGACRYSTIC GRANITE SAMPLE 102-3A 39 40 40 39 40 39 Ar/ Ar Ar/ Ar r Ca/K % Atm % Ar Ar*/ K 0.003253 ± 0.000184 0.005280 ± 0.000432 0.026 0.000 96.12 0.05 7.315 ± 10.268 0.003037 ± 0.000144 0.007950 ± 0.000387 0.012 0.000 89.71 0.09 12.913 ± 5.403 0.001804 ± 0.000324 0.023044 ± 0.001042 0.004 0.000 53.28 0.08 20.261 ± 4.269 0.002041 ± 0.000185 0.017804 ± 0.000586 0.009 0.000 60.29 0.13 22.288 ± 3.159 0.001352 ± 0.000144 0.023573 ± 0.000627 0.005 0.000 39.91 0.19 25.477 ± 1.940
Age 93.22 ± 127.54 161.45 ± 64.62 247.26 ± 48.68 270.23 ± 35.57 305.79 ± 21.42
2.25 2.50 <2.75 <3.00 <3.50
0.000790 ± 0.000102 0.000407 ± 0.000085 0.000192 ± 0.000062 0.000317 ± 0.000037 0.000170 ± 0.000012
0.028341 ± 0.000562 0.031525 ± 0.000527 0.032267 ± 0.000406 0.031507 ± 0.000314 0.032491 ± 0.000210
0.004 0.004 0.002 0.002 0.003
0.000 0.000 0.000 0.000 0.000
23.32 12.02 5.68 9.36 5.03
0.32 0.40 0.68 1.13 4.01
27.051 ± 1.200 27.904 ± 0.929 29.228 ± 0.681 28.767 ± 0.451 29.228 ± 0.220
323.08 ± 13.13 332.38 ± 10.11 346.74 ± 7.35 341.74 ± 4.89 346.73 ± 2.37
<4.00 <4.50 <5.50 <6.50 <7.50
0.000098 ± 0.000011 0.000151 ± 0.000013 0.000091 ± 0.000010 0.000054 ± 0.000008 0.000020 ± 0.000033
0.033120 ± 0.000205 0.032661 ± 0.000195 0.033111 ± 0.000197 0.033667 ± 0.000210 0.033471 ± 0.000358
0.002 0.002 0.001 0.001 0.001
0.002 0.001 0.003 0.003 0.008
2.88 4.45 2.67 1.60 0.59
4.70 3.83 10.99 21.18 17.81
29.323 ± 0.210 29.254 ± 0.212 29.393 ± 0.196 29.227 ± 0.196 29.700 ± 0.431
347.76 ± 2.27 347.01 ± 2.29 348.51 ± 2.12 346.72 ± 2.12 351.81 ± 4.64
<7.80
0.000018 ± 0.000015
0.033230 ± 0.000241
0.001
0.067
0.53
34.41
29.933 ± 0.257
354.33 ± 2.77 Continued
36
40
39
40
Power 0.75 1.25 1.50 1.75 2.00
TABLE 7. Ar/ Ar ANALYSES OF MUSCOVITE IN MEGACRYSTIC GRANITE SAMPLE 102-3A (continued) 40 39 38 37 36 40 Ar Ar Ar Ar Ar Blank Ar 2.594 ± 0.009 0.015 ± 0.001 0.003 ± 0.001 0.001 ± 0.001 0.009 ± 0.000 0.035 3.272 ± 0.010 0.027 ± 0.001 0.003 ± 0.001 0.001 ± 0.001 0.011 ± 0.000 0.034 1.006 ± 0.004 0.024 ± 0.001 0.001 ± 0.001 0.001 ± 0.001 0.002 ± 0.000 0.034 2.024 ± 0.009 0.037 ± 0.001 0.002 ± 0.001 0.001 ± 0.001 0.005 ± 0.000 0.034 2.367 ± 0.007 0.057 ± 0.001 0.002 ± 0.001 0.001 ± 0.001 0.004 ± 0.000 0.034
2.25 2.50 <2.75 <3.00 <3.50
3.242 ± 0.010 3.631 ± 0.012 5.983 ± 0.014 10.203 ± 0.016 34.953 ± 0.036
0.093 ± 0.002 0.115 ± 0.002 0.194 ± 0.002 0.323 ± 0.003 1.140 ± 0.007
0.003 ± 0.001 0.002 ± 0.001 0.003 ± 0.001 0.006 ± 0.001 0.016 ± 0.001
0.001 ± 0.001 0.001 ± 0.001 0.001 ± 0.001 0.001 ± 0.001 0.001 ± 0.001
0.003 ± 0.000 0.002 ± 0.000 0.002 ± 0.000 0.004 ± 0.000 0.006 ± 0.000
0.034 0.035 0.034 0.035 0.034
Atmos 40/36 289.649 289.649 289.649 289.649 289.649 289.649 289.649 289.649 289.649 289.649
<4.00 40.154 ± 0.045 1.335 ± 0.008 0.019 ± 0.001 0.002 ± 0.001 0.005 ± 0.000 0.033 289.649 <4.50 33.194 ± 0.032 1.089 ± 0.006 0.016 ± 0.001 0.002 ± 0.001 0.006 ± 0.000 0.033 289.649 <5.50 93.877 ± 0.075 3.120 ± 0.018 0.042 ± 0.003 0.003 ± 0.002 0.009 ± 0.001 0.034 289.649 <6.50 177.990 ± 0.171 6.015 ± 0.037 0.079 ± 0.005 0.005 ± 0.003 0.010 ± 0.001 0.034 289.649 39 Note: AOR-579: 102-3A Ms 40/60. J value: 0.007250 ± 0.000024. Mass: 35.0 mg; Approx. 4.30% K, 0.11% Ca. Volume K: 28.25 × 1E-10 3 cm NTP. Integrated age: 349.79 ± 1.75 Ma. Initial 40/36: 226.49 ± 27.68 (mean square of weighted deviates [MSWD] = 5.18, isochron between 39 0.62 and 1.76). Correlation age: 350.39 ± 2.19 Ma (100.0% of Ar, steps marked by >). MSWD: 4.241. Plateau age: 350.50 ± 1.76 Ma (98.7% 39 of Ar, steps marked by <). Modified error: 1.85.
corresponding 40Ar/39Ar ages (e.g., Boundy et al., 1997) and was a problem with other minerals analyzed from the Asis Lithodeme, this muscovite age appears to be unaffected by excess argon, because the initial 40Ar/36Ar is similar to present-day atmospheric values. Thus, we infer that it dates cooling through ~350 °C using the actual grain size and experimental data from Hames and Bowring (1994). PRESSURE-TEMPERATURE-TIME (P-T-T) DATA The P-T estimates have been combined with the geochronology and the structural sequence in a P-T-t graph (Fig. 14). This shows a clockwise path that is most complete for the amphibolites: eclogite through amphibolite, greenschist, and sub-greenschist facies. Only the latter part of this path can be documented for the metasedimentary and granitoid rocks. However, the prob-
ability that the amphibolites were originally either lavas or dikes in the metasedimentary rocks and the presence of metasedimentary xenoliths with a foliation parallel to that in the host granitoid rocks suggest that they experienced a similar structural history. On the other hand, the 351 ± 2 Ma muscovite cooling age from the granite is slightly older than the 346 ± 3 Ma U-Pb zircon age from the amphibolite, suggesting that the megacrystic granite rose through ~350 °C before the metasedimentary and amphibolitic rocks. This is consistent with the kinematic indicators that suggest the granite moved southwestward relative to the surrounding rocks and is probably reflected in the mylonized borders of the granite (Fig. 3). Relative motion between the granite and the country rocks probably took place along the shear zones between them during greenschist-facies deformation (Fig. 3). Thus, a combined P-T-t path for the amphibolites and metasedimentary rocks indicates that they passed through the following facies:
P-T-t constraints on exhumation following subduction in the Rheic Ocean
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351 ± 2 Ma
De 346±3 Ma
Piaxtla and Mimilulco Acatlán Complex Meza-Figueroa et al., 2003
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1. The eclogite facies based on omphacite analyses associated with fabric Se; this is inferred to have taken place at 346 ± 3 Ma. Unfortunately, the eclogite-facies temperatures and pressures are constrained by only two analyses with wide limits: an average of the lowest values gives ~650–750 °C at ≥14 kilobars. 2. The upper amphibolite facies during the Mississippian. The absence of minerals typical of the granulite facies suggests that the ~750–850 °C temperatures estimated for the amphiboles, magnesio-taramite and barroisite, are too high; however, the presence of migmatitic leucosomes in all rock units indicates temperatures above the minimum melting curve, i.e., upper amphibolite facies and >600 °C. 3. The lower amphibolite to greenschist facies based on the presence of epidote and actinolite. The age of these fabrics is presently unknown, but the structural fabrics are congruent with earlier fabrics and so probably represent a continuation of the Mississippian deformation. 4. The sub-greenschist facies associated with the crenulation cleavages and late folds. The age of the crenulation cleavage fabrics is presently unconstrained. Permo-Triassic and Cretaceous–Eocene (Laramide) events also affected the region (Keppie et al., 2004a, 2006). The Mississippian age for exhumation of the high-grade metamorphism in the Asis Lithodeme is considerably younger than the Late Ordovician age previously assumed (Ortega-Gutiérrez et al., 1999; Fig. 2), and provides important constraints on tectonic interpretations. For example, the interpretation that the megacrystic granites were decompression melts following eclogite-facies metamorphism cannot be retained given the Ordovician age of these granitoids (Sánchez-Zavala et al., 2004; Talavera-Mendoza et al., 2005; Murphy et al., 2006). On the other hand, migmatization produced by decompression melting of eclogitic rocks is supported by the geochronology. Rapid exhumation to the surface is suggested by outcrops 10 km to the north of the Asis area, where eclogitic rocks are unconformably overlain by the unmetamorphosed, latest Upper Fammenian Otates Formation (Vachard et al., 2000; Vachard and Flores de Dios, 2002; Elías-Herrera et
8
lite
Figure 13. Ar/ Ar incremental release spectrum from muscovite in megacrystic granite sample 102–3A.
ib o
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ph
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Ca. 351 Ma 40 Ar/ 39Ar muscovite in granite
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Asis Lithodeme 40 This paper
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ne oxe Pyr nfels Hor 700
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Temperature (°C) Figure 14. Pressure-temperature (P-T) graph showing ages and structural events from the Asis Lithodeme, Piaxtla Suite, southern México. The P-T path from Meza-Figueroa et al. (2003) is shown for comparison.
al., 2004; Derycke-Khatir et al., 2005), i.e., dated ca. 365 Ma. This area is adjacent to Mimilulco, one of the areas studied by Meza-Figueroa et al. (2003), where the P-T path reveals lower temperatures as the rocks passed through the amphibolite facies, but they are similar for temperatures below ~500 °C (Fig. 14). Unfortunately, geochronological data allowing us to compare the relative ages of these rocks are presently not available; however, the unconformity beneath the Patlanoaya Formation indicates the likelihood of diachronism between the two areas (Fig. 2). DISCUSSION Eclogite-facies rocks are typical of regions that have undergone subduction (Ernst, 1988). In the Asis area, a younger limit for the eclogite-facies metamorphism is indicated by the concordant U-Pb zircon age of 346 ± 3 Ma, and this metamorphism was quickly followed by decompression melting producing migmatites at ca. 348–330 Ma. Although age data are not available for the last stage of exhumation, it may also have been rapid. Deposition of unmetamorphosed sedimentary rocks of the latest Upper Fammenian Otates Formation on eclogitic rocks (Elías-Herrera et al., 2004) in an area 10 km to the north of the Asis area suggests either synchronous basin subsidence and exhumation in adjacent areas or diachronous exhumation. Rapid exhumation of eclogites like that recorded in the Asis area has generally been related to
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continent-continent collision. Thus Ernst (1988) relates rapid exhumation of blueschists to subduction of a buoyant mass during continental collision. Hynes (2002) infers extrusion bounded below by thrusts and above by normal faults as the exhumation mechanism during continental collision, and Reddy et al. (2003) suggest extensional exhumation in areas of transtension during continental collision. Although the record of shallowly plunging stretching lineations in the Asis Lithodeme is compatible with any of these models, the presence of depositional basins (the Otates and Patlanoaya Formations) close to the rapidly exhumed Asis Lithodeme favors a combination of the second and third models, i.e., extrusion
Tehuitcingo ophiolite
480-460 Ma COSOLTEPEC FM plutons (rise prism)
and extensional exhumation during oblique continental collision. Further study of the boundaries of the eclogitic rocks is required to resolve this issue. Continent-continent collision is consistent with early Carboniferous paleogeographic reconstructions indicating that the amalgamation of Pangea took place by subduction of the northern margin of Gondwana (including México) beneath remnants of the Rheic Ocean, followed by oblique dextral collision between Gondwana and Laurentia, with the Acatlán Complex of southern México caught in the collision zone (Figs. 15 and 16). The Late Devonian–Mississippian age of the subduction-related metamorphism, although incompatible with an origin in the Iapetus Ocean (Ortega-Gutiérrez et al., 1999), which had closed by Silurian times (Keppie, 1993; Hibbard et al., 2002), is consistent with an origin in the Rheic Ocean (Keppie and Ramos, (1999). ACKNOWLEDGMENTS
Shelf
Rheic Ocean Suture
GONDWANA (OAXACAN CX)
SOUTHERN LAURENTIA
Asis Lithodeme PIAXTLA SUITE Rift + Rise
LATE DEVONIANMISSISSIPPIAN Figure 15. Tectonic cross-section showing subduction of the Asis Lithodeme in the Devono-Carboniferous.
We are grateful to Drs. R.A. Strachan and Z. de Cserna for their constructive reviews. We would like to acknowledge PAPIIT (Programa de Apoyo a Proyectos de Investigación e Innovación Tecnológica) grant IN103003 to JDK, National Science Foundation grants EAR 0308105 to RDN and EAR 0308437 to BVM, and a NSERC (National Science and Engineering Research Council) Discovery grant to JD and JBM. We would also like to thank Miguel Morales for assistance with drawing the figures. This paper represents a contribution to IGCP (International Geological Correlation Program) Project 498, The Rheic Ocean. REFERENCES CITED
BALTICA
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370 Ma Figure 16. Late Devonian reconstruction showing the Acatlán Complex of southern México sandwiched between Gondwana and Laurentia during the amalgamation of Pangea (modified after Keppie, 2004).
Berman, R.G., 1991, Thermobarometry using multi-equilibrium calculations: A new technique, with petrological applications: Canadian Mineralogist, v. 29, p. 833–855. Boundy, T.M., Hall, C.M., Li, G., Essene, E.J., and Halliday, A.N., 1997, Finescale isotopic heterogeneities and fluids in the deep crust: A 40Ar/39Ar laser ablation and TEM study of muscovites from a granulite–eclogite transition zone: Earth and Planetary Science Letters, v. 148, p. 223–242, doi: 10.1016/S0012-821X(97)00036-8. Brown, E.H., 1977, The crossite content of Ca-amphibole as a guide to pressure of metamorphism: Journal of Petrology, v. 18, p. 53–72. Carswell, D.A., 1990, Eclogite facies rocks: New York, Chapman & Hall. Clark, A.H., Archibald, D.A., Lee, A.W., Farrar, E., and Hodgson, C.J., 1998, Laser Probe 40Ar/39Ar ages of early- and late-stage alteration assemblages, Rosario porphyry copper-molybdenum deposit, Collahuasi District, I Region, Chile: Economic Geology and the Bulletin of the Society of Economic Geologists, v. 93, p. 326–337. Coleman, R.G., Lee, D.E., Beatty, L.B., and Brannock, W.W., 1965, Eclogites and eclogites: Their differences and similarities: Geological Society of America Bulletin, v. 76, p. 483–508. DeGraaff-Surpless, K., Graham, S.A., Wooden, J.L., and McWilliams, M.O., 2002, Detrital zircon provenance analysis of the Great Valley Group, California: Evolution of an arc-forearc system: Geological Society of America Bulletin, v. 114, p. 1564–1580, doi: 10.1130/0016-7606(2002)114<1564: DZPAOT>2.0.CO;2. Derycke-Khatir, C., Vachard, D., Dégardin, J.-M., Flores de Dios, A., Buitrón, B., and Hansen, M., 2005, Late Pennsylvanian and Early Permian chondichthyan microremains from San Salvador Patlanoaya (Puebla, Mexico): Geobios, v. 38, p. 43–55, doi: 10.1016/j.geobios.2003.06.008. Elías-Herrera, M., and Ortega-Gutiérrez, F., 2002, Caltepec fault zone—An Early Permian dextral transpressional boundary between the Proterozoic Oaxacan and Paleozoic Acatlán complex, southern Mexico, and regional implications: Tectonics, v. 21, p. 1–19, doi: 10.1029/2000TC001278.
P-T-t constraints on exhumation following subduction in the Rheic Ocean Elías-Herrera, M., Ortega-Gutiérrez, F., Sánchez-Zavala, J.L., Reyes-Salas, A.M., Mácias-Romo, C., and Irriondo, A., 2004, New geochronological and stratigraphic data related to the Paleozoic evolution of the high-pressure Piaxtla Group, Acatlán Complex, southern Mexico: Libro de Resumenes, IV Reunión Nacional de Ciencias de la Teirra, Juriquilla, Querétar, p. 150. Ellis, D.J., and Green, D.H., 1979, An experimental study of the effect of Ca upon garnet-clinopyroxene Fe-Mg exchange equilibria: Contributions to Mineralogy and Petrology, v. 71, p. 13–22, doi: 10.1007/BF00371878. Ernst, W.G., 1988, Tectonic history of subduction zones inferred from retrograde blueschist P-T paths: Geology, v. 16, p. 1081–1084, doi: 10.1130/00917613(1988)016<1081:THOSZI>2.3.CO;2. Ganguly, J., 1979, Garnet and clinopyroxene solid solutions, and geothermometry based on Fe-Mg distribution coefficient: Geochimica et Cosmochimica Acta, v. 43, p. 1021–1029, doi: 10.1016/0016-7037(79)90091-7. Hames, W.E., and Bowring, S.A., 1994, An empirical evaluation of the argon diffusion geometry in muscovite: Earth and Planetary Science Letters, v. 124, p. 161–167, doi: 10.1016/0012-821X(94)00079-4. Heaman, L., and Parrish, R.R., 1991, U-Pb geochronology of accessory minerals: Mineralogical Association of Canada Short Course Handbook, v. 19, p. 59–102. Hibbard, J.P., Stoddard, E.F., Stoddard, E.F., Secor, D.T., and Dennis, A.J., 2002, The Carolina Zone: Overview of Neoproterozoic to Early Paleozoic peri-Gondwanan terranes along the eastern flank of the southern Appalachians: EarthScience Reviews, v. 57, p. 299–339, doi: 10.1016/S0012-8252(01)00079-4. Holland, T.J.B., 1980, The reaction albite = jadeite + quartz determined experimentally in the range 600–1200 °C: The American Mineralogist, v. 65, p. 129–134. Holland, T.J.B., and Powell, R., 1990, An enlarged and updated internally consistent thermodynamic dataset with uncertainties and correlations, the system K2O-Na2O-CaO-MgO-FeO-Fe2O3-Al2O3-TiO2-SiO2-CH2O2: Journal of Metamorphic Geology, v. 8, p. 89–124. Hynes, A., 2002, Encouraging the extrusion of deep-crustal rocks in collision zones: Mineralogical Magazine, v. 66, no. 1, p. 5–24, doi: 10.1180/00264 61026610013. Hynes, A., and Forest, R.C., 1988, Empirical garnet-muscovite geothermometry in low-grade metapelites, Selwyn Range Canadian Rockies: Journal of Metamorphic Geology, v. 6, p. 297–309. Keppie, J.D., 2004, Terranes of Mexico revisited: A 1.3 billion year odyssey: International Geology Review, v. 46, p. 765–794. Keppie, J.D., and Ramos, V.S., 1999, Odyssey of terranes in the Iapetus and Rheic Oceans during the Paleozoic, in Ramos, V.A., and Keppie, J.D., eds., Laurentia-Gondwana connections before Pangea: Boulder, Colorado, Geological Society of America Special Paper 336, p. 267–276. Keppie, J.D., Sandberg, C.A., Miller, B.V., Sánchez-Zavala, J.L., Nance, R.D., and Poole, F.G., 2004a, Implications of latest Pennsylvanian to Middle Permian paleontological and U-Pb SHRIMP data from the Tecomate Formation to re-dating tectonothermal events in the Acatlán Complex, southern Mexico: International Geology Review, v. 46, p. 745–753. Keppie, J.D., Nance, R.D., Powell, J.T., Mumma, S.A., Dostal, J., Fox, D., Muise, J., Ortega-Rivera, A., Miller, B.V., and Lee, J.W.K., 2004b, Mid-Jurassic tectonothermal event superposed on a Paleozoic geological record in the Acatlán Complex of southern Mexico: Hotspot activity during the breakup of Pangea: Gondwana Research, v. 7, p. 238–260, doi: 10.1016/S1342-937X(05)70323-3. Keppie, J.D., Nance, R.D., Fernández-Suarez, J., Storey, C.D., Jeffries, T.E., and Murphy, J.B., 2006, Detrital zircon data from the eastern Mixteca terrane, southern Mexico: Evidence for an Ordovician–Mississippian continental rise and a Permo-Triassic clastic wedge adjacent to Oaxaquia: International Geology Review, v. 46, no. 2., p. 97–111. Krogh, E.J., and Raheim, A., 1978, Temperature and pressure dependence of Fe-Mg partitioning between garnet and phengite, with particular reference to eclogites: Contributions to Mineralogy and Petrology, v. 66, p. 75–80, doi: 10.1007/BF00376087. Ludwig, K.R., 1988, PBDAT for MS-DOS: A computer program for IBM-PC compatibles for processing raw Pb-U-Th isotope data: Reston, Virginia, U.S. Geological Survey Open-File Report 88-0542. Ludwig, K.R., 2001, SQUID 1.02: A user’s manual: Berkeley, Berkeley Geochronology Center Special Publication 2, 19 p. Ludwig, K.R., 2003, Isoplot/Ex, version 2.49, a geochronological toolkit for Microsoft Excel: Berkeley, California, Berkeley Geochronological Center Special Publication 1a. Malone, J.W., Nance, R.D., Keppie, J.D., and Dostal, J., 2002, Deformational history of part of the Acatlán Complex—Late Ordovician–Early Silurian and
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Early Permian orogenesis in southern Mexico: Journal of South American Earth Sciences, v. 15, p. 511–524, doi: 10.1016/S0895-9811(02)00080-9. Massonne, H.J., and Schreyer, W., 1987, Phengite geobarometry based on the limiting assemblage with K-feldspar, phlogopite, and quartz: Contributions to Mineralogy and Petrology, v. 96, p. 212–224. Meza-Figueroa, D., Ruiz, J., Talavera-Mendoza, O., and Ortega-Gutiérrez, F., 2003, Tectonometamorphic evolution of the Acatlán complex eclogites (southern Mexico): Canadian Journal of Earth Sciences, v. 40, p. 27–44, doi: 10.1139/e02-093. Murphy, J.B., Keppie, J.D., Nance, R.D., Miller, B.V., Dostal, J., Middleton, M., Fernandez-Suárez, J., and Jeffries, T.E., 2006, Geochemistry and U-Pb protolith ages of eclogitic rocks of the Asis Lithodeme, Piaxtla Suite, Acatlán Complex, southern Mexico: Cambro-Ordovician rift–passive margin sequence cut by Ordovician rift or arc granitoids on the southern margin of the Rheic Ocean: Journal of the Geological Society, London, v. 163, p. 683–695. Newton, R.C., and Perkins, D., 1982, Thermodynamic calibration of geobarometers based on the assemblage garnet-plagioclase-orthopyroxene (clinopyroxene)-quartz: American Mineralogist, v. 67, p. 203–222. North American Commission on Stratigraphic Nomenclature, 1983, North American Stratigraphic Code: American Association of Petroleum Geologists Bulletin, v. 67, p. 841–875. Ortega-Gutiérrez, F., Eliás-Herrera, M., Reyes-Salas, M., Maciás-Romo, C., and Lopez, R., 1999, Late Ordovician–Early Silurian continental collision orogeny in southern Mexico and its bearing on Gondwana-Laurentia connections: Geology, v. 27, p. 719–722, doi: 10.1130/00917613(1999)027<0719:LOESCC>2.3.CO;2. Raheim, A., and Green, D.H., 1974, Experimental determination of the temperature and pressure dependence of the Fe-Mg partition coefficient for coexisting garnet and clinopyroxene: Contributions to Mineralogy and Petrology, v. 48, p. 179–203, doi: 10.1007/BF00383355. Ratajeski, K., Glazner, A.F., and Miller, B.V., 2001, Geology and geochemistry of mafic to felsic plutonic rocks in the Cretaceous intrusive suite of Yosemite Valley, California: Geological Society of America Bulletin, v. 113, p. 1486– 1502, doi: 10.1130/0016-7606(2001)113<1486:GAGOMT>2.0.CO;2. Reddy, S.M., Wheeler, J., Butler, R.W.H., Cliff, R.A., Freeman, S., Inger, S., Pickles, C., and Kelley, S.P., 2003, Kinematic reworking and exhumation within the convergent Alpine Orogen: Tectonophysics, v. 365, p. 77–102, doi: 10.1016/S0040-1951(03)00017-9. Ryburn, R.J., Raheim, A., and Green, D.H., 1976, Determination of the P-T paths of natural eclogites during metamorphism, a record of subduction: Lithos, v. 8, p. 317–328. Sánchez-Zavala, J.L., Ortega-Gutiérrez, F., Keppie, J.D., Jenner, A.G., Belousova, E., and Maciás-Romo, C., 2004, Ordovician and Mesoproterozoic zircons from the Tecomate formation and Esperanza Granitoids, Acatlán Complex, southern Mexico: Local provenance in the Acatlán and Oaxacan complexes: International Geology Review, v. 46, p. 1005–1021. Spear, F.S., 1980, NaSi-CaAI exchange equilibrium between plagioclase and amphibole: Contributions to Mineralogy and Petology, v. 72, p. 33–41. Talavera-Mendoza, O., Ruíz, J., Gehrels, G.E., Meza-Figueroa, D.M., VegaGranillo, R., and Campa-Uranda, M.F., 2005, U-Pb geochronology of the Acatlán Complex and implications for the Paleozoic paleogeography and tectonic evolution of southern Mexico: Earth and Planetary Science Letters, v. 235, p. 682–699. Vachard, D., and Flores de Dios, A., 2002, Discovery of latest Devonian / earliest Mississippian microfossils in San Salvador Patlanoaya (Puebla, Mexico): Biogeographic and geodynamic consequences: Comptes Rendus Geoscience, v. 334, p. 1095–1101, doi: 10.1016/S1631-0713(02)01851-5. Vachard, D., Flores de Dios, A., Buitrón, B.E., and Grahales, M., 2000, Biostratigraphie par fusilines des calcaires Carbonifères et Permienes de San Salvador Patlanoaya (Puebla, Mexique): Geobios, v. 33, p. 5–33, doi: 10.1016/S0016-6995(00)80145-X. Waters, D., and Martin, H.N., 1993, Geobarometry in phengite-bearing eclogites: Terra Abstracts (European Union of Geosciences), v. 5, p. 410–411. Yañez, P., Ruiz, J., Patchett, P.J., Ortega-Gutiérrez, F., and Gehrels, G.E., 1991, Isotopic studies of the Acatlán complex, southern Mexico: Implications for Paleozoic North American tectonics: Geological Society of America Bulletin, v. 103, p. 817–828. Yassaghi, A., 2000, Deformational behavior of quartz and feldspar in quartzites within shear zones in the Adelaide Hills, South Australia: Iranian International Journal of Science, v. 1, p. 109–130. MANUSCRIPT ACCEPTED BY THE SOCIETY 3 OCTOBER 2006 Printed in the USA
Geological Society of America Special Paper 423 2007
Life and death of a Cambrian–Ordovician basin: An Andean threeact play featuring Gondwana and the Arequipa-Antofalla terrane Sven O. Egenhoff* Colorado State University, Department of Geosciences, 322 Natural Resources Building, Fort Collins, Colorado 80523-1482, USA
ABSTRACT Reconstruction of the facies architecture and geometries of the Cambrian–Ordovician succession in the Central Andes of southern Bolivia and northwestern Argentina reveals a tripartite basin history that closely corresponds to interpretations of regional plate tectonic movements. The analysis of basin deposits enabled tracing and timing of movements of the Arequipa-Antofalla terrane, which initiated, fed, and terminated a basin between the terrane and Gondwana during the early Paleozoic. Tectonic movements started in the Cambrian and led to the formation of an extensional basin. Stretching was more pronounced in southern Bolivia than in northwestern Argentina, resulting in widening of the basin to the north. This was produced by a counterclockwise rotation of the Arequipa-Antofalla terrane relative to Gondwana, with a Euler pole in northwest Argentina. Tectonic movements reversed in the late Early Ordovician (Expansograptus holmi biozone), with the terrane rotating clockwise back against Gondwana. Consequently, the extensional basin turned into a foreland trough, with its western part undergoing high subsidence. A forebulge developed on the eastern shelf, which triggered westward progradation of a delta, thereby significantly reducing basin width. Although compressional movements advanced from west to east, reliable biostratigraphic calibration was established only for facies on the distal part of the overridden plate as graptolite ecology and abundance in this area enhanced biostratigraphic resolution. In the Late Ordovician, the basin closed in northwestern Argentina but continued to accumulate coarse-grained, partially glacigenic debris in southern Bolivia. Keywords: Argentina, Bolivia, graptolite, terrane, basin fill INTRODUCTION During the Cambrian–Ordovician, the western rim of Gondwana was an active continental margin that recorded the accretion of exotic and para-autochthonous terranes (Bahlburg and Hervé, 1997; Pankhurst and Rapela, 1998). One of the more prominent peri-Gondwanan microcontinents that today forms surface and *E-mail:
[email protected].
subsurface rocks in Peru, Bolivia, Argentina, and Chile is the Arequipa-Antofalla terrane (Gohrbandt, 1992; Forsythe et al., 1993; Bahlburg and Hervé, 1997; Loewy et al., 2004). Paleomagnetic data suggest a clockwise rotation of this terrane away from Gondwana and the concomitant formation of a Cambrian–Ordovician basin, followed by a counterclockwise backrotation and reaccretion of the terrane during the early Paleozoic (Forsythe et al., 1993). However, a matter still debated (Bahlburg, 1990, 1991; Bahlburg and Hervé, 1997; Coira et al., 1999) is the time at which
Egenhoff, S.O., 2007, Life and death of a Cambrian–Ordovician basin: An Andean three-act play featuring Gondwana and the Arequipa-Antofalla terrane, 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. 511–524, doi: 10.1130/2007.2423(26). For permission to copy, contact
[email protected]. ©2007 Geological Society of America. All rights reserved.
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512
Egenhoff
extension switched to compression that eventually resulted in the accretion of the terrane to the continent by the end of the Ordovician (Forsythe et al., 1993; Bahlburg and Hervé, 1997). This study addresses these issues by using the basin fill between Gondwana and the Arequipa-Antofalla terrane to reconstruct the relative movements of this terrane during the Cambrian–Ordovician. Analysis of sediment geometries coupled with a detailed biostratigraphic framework enables precise relative dating of profound changes in the evolution of this basin, and therewith of the terrane-Gondwana interaction. New data are presented from southern Bolivia, an area whose early Paleozoic paleogeography has been largely ignored (Gohrbandt, 1992; Sempere, 1995), and the Bolivian basin evolution is compared with time-equivalent reconstructions from adjacent northwestern Argentina (Moya, 1988; Moya et al., 1994; Buatois and Mángano, 2003). This study demonstrates that the integration of Bolivian and Argentinean data from the Cambrian–Ordovician succession is crucial for understanding the differential basin evolution along a north-south transect parallel to the basin axis, and therewith the relative movements of the Arequipa-Antofalla terrane.
The Ordovician basin history in the southern Central Andes is defined by three episodes: basin opening at some time in the Cambrian; the reversal in the movement of the Arequipa-Antofalla terrane in the late Early Ordovician, causing a change from an extensional to a compressional regime; and basin closure in the latest Ordovician. STRATIGRAPHY Cambrian–Ordovician sedimentary rocks make up the bulk of the surface exposures in the center of southern Bolivia. The region is subdivided into three main zones bound by major Andean overthrusts (Fig. 1; Kley, 1993) with rocks that are younger from east to west. The oriental zone, comprising the eastern portion of the Eastern Cordillera and the Subandean belt (Fig. 1; Reutter et al., 1988; Kley, 1993), is composed of Cambrian to Lower Ordovician rocks, with older strata more common in the east. The central zone, where only Middle Ordovician rocks crop out, is 10–20 km wide and runs through the city of Tupiza (Fig. 1). Upper Ordovician rocks form the largest portion of the western part of the Eastern Cordillera and occur in scattered outcrops in
10°
Key
BRAZIL
N
BOLIVIA
Silurian to Tertiary
Cochabamba
Middle Ordovician
PERU 15°
Upper Ordovician
La Paz Sta.Cruz
Lower Ordovician
PACIFIC OCEAN
Tica Tica Potosì
CHILE
Tarija
20°
Cambrian studied localities
PARAGUAY
city ARGENTINA 65°
70°
25° 60°
50 km Rio Palomita
BOLIVIA
Challa Mayu
Rio Marquina Soniquera
Tupiza Mauka Huasi
San Pablo de Lipez
Pilar Punta Loma Chaupinuo Cieneguillas Abra Negra Taraya
Mal Paso
N Sella
Sama
Tarija Pinos Tajsara
22°
67°
Incahuasi
Culpina
66°
ARGENTINA
22°
65°
Figure 1. Geological map of the study area in southern Bolivia, focusing on the lower Paleozoic succession (based on Reutter et al., 1994). Positions of investigated localities are indicated by white stars. The Cambrian–Ordovician rocks show an overall younging to the west.
Life and death of a Cambrian–Ordovician basin: An Andean three-act play the Altiplano region. Stratigraphic and sedimentary facies data are summarized in Figures 2–9. Stratigraphic nomenclature for the Ordovician follows Finney (2005). BASIN FORMATION AND EARLY INFILL The Cambrian succession has been studied in the localities of Pinos, ~25 km southwest of Tarija, Tajsara, and Sama (Fig. 1). A generalized section together with descriptions and sedimentological interpretations of the four stratigraphic units—the Camacho, Torohuayco, Sama, and Iscayachi Formations—which are completely or partially Cambrian, is shown in Figure 2. Table 1 illustrates how these formations relate to their northwest Argentinean counterparts (cf. Kumpa and Sanchez, 1988; Moya, 1988; Buatois and Mángano, 2003). The entire Cambrian to Lower Ordovician basin fill, composed of continental to deep-shelf strata, is an overall transgressive succession in which facies belts shift successively onto the Brazilian Shield to the east and the basin widens. The conglomerates and sandstones of the Camacho Formation that form the earliest basin fill represent fluvial facies deposited from braided rivers (Fig. 2). The coarse grain sizes argue for an overall high relief of the depositional site and the adjacent provenance area and confirm the position of the Tarija region in a rather marginal part of the basin during the initial basin-forming phase in the Cambrian (Gohrbandt, 1992). The abundance of marine trace fossils throughout the Torohuayco Formation indicates that the basin had already become entirely marine in the Cambrian. The overlying Sama and Iscayachi Formations represent more open marine conditions. The presence of channels and delta front sands in the Iscayachi Formation indicates the development of a delta system in the Late Cambrian. The mud-dominated turbiditic Obispo and Taraya Formations overlie the Iscayachi delta and record another deepening of the depositional area in the Early Ordovician. The entire Cambrian succession is ~2500 m thick in the Sama area; the cumulative thickness of the Camacho, Torohuayco, and Sama Formations is ~1800 m (Kley, 1993), and ~700 m of the 1000-m thick Iscayachi Formation are also Cambrian (Egenhoff, 2000). Kley (1993) reported that the Cambrian strata thicken toward the south. However, this trend does not continue into northwestern Argentina, where most published data indicate lesser thicknesses for Cambrian strata (combined data from Kumpa and Sanchez, 1988, and Buatois and Mángano, 2003: 420–1150 m) compared with southern Bolivia (1800 m minimum thickness; Kley, 1993). The thickening of Cambrian strata to the south within southern Bolivia is interpreted as a local phenomenon, and thickness variations indicate a basin with laterally strongly varying subsidence patterns that may be controlled by either tectonics or compaction or both. The increase of Cambrian sediment thicknesses between Sama (southern Bolivia; Fig. 1) and the BolivianArgentinean border may reflect higher sediment input from the Cambrian Iscayachi delta (Egenhoff, 2000) that delivered more siliciclastic material to this particular basin area. The southward
513
decrease of sediment thickness into northwestern Argentina may reflect a gradual shift from the delta, with very high clastic input into an adjacent shoreline accumulating less sediment (cf. Buatois and Mángano, 2003). This decrease in thickness of Cambrian strata may represent overall reduced subsidence in the Argentinean part of the basin. A comparison of thicknesses of Lower Ordovician strata in Argentina (2200 m; Moya, 1988) and southern Bolivia (4200 m minimum; Egenhoff, 2000) indicates that the differential subsidence pattern continued into the Early Ordovician. It may have been related to differences in extensional tectonism, which likely was more pronounced in southern Bolivia than in northwestern Argentina. Consequently, the basin was opening more rapidly in southern Bolivia, and the Cambrian marine ingression likely flooded the basin from the north. All subsequent transgressions have most probably advanced from north to south within the basin. The succession of northwestern Argentina is characterized by abundant hiatuses within the Cambrian and Lower Ordovician interval (Buatois and Mángano, 2003), whereas the southern Bolivian sedimentary package is not only thicker, but also biostratigraphically more complete (Egenhoff et al., 2004). The general shallowing and narrowing of the basin toward the south are further reflected in the facies diachrony of Argentinean and Bolivian strata. Southern Bolivia experienced a major transgression in the Lower Tremadocian Rhabdinopora flabeliformis biozone, where the strata were deposited below storm wave base. In northwestern Argentina this transgression is recorded in the slightly higher Anisograptus sp. biozone (Moya, 1988, 2002). CHANGE FROM EXTENSIONAL TO COMPRESSIONAL REGIME Twelve Lower Ordovician shelf sections have been measured along a transect between Sella in the east and Challa Mayu in the west (Fig. 1). Correlation of the sections is based on their graptolite biostratigraphy (Egenhoff et al., 2004; Fig. 3). The succession is subdivided into the Obispo, Taraya, Pircancha, Sella, and Rumi Orkho Formations (Rivas et al., 1969; Egenhoff, 2000; Schönian, 2000), with their stratigraphic relationships shown in Figures 3 and 4. A summary of the facies and environmental interpretations is presented in Figure 5. After a major transgression in the Early Tremadocian (Egenhoff, 2000), the Lower Ordovician shelf sediments of the Obispo and Taraya Formations, exposed throughout the study area, record turbidite deposition on a submarine ramp (Fig. 6). This turbidite system was fed by a delta located farther to the east on the Brazilian Shield. That a delta served as sediment source for the Bolivian shelf is inferred from the enormous thickness of the Obispo Formation and the interfingering of the delta facies of the Obispo Formation with turbidite ramp sediments of the Pircancha Formation in the upper Lower Ordovician (discussed later). The Taraya Formation consists exclusively of pelitic strata, in contrast to the Obispo Formation, which contains numerous turbidite sand beds. It is confined to an area a few kilometers wide
514
Egenhoff
Figure 2. Generalized section with facies description and sedimentological interpretation of the Cambrian succession in southern Bolivia that comprises the Camacho, Torohuayco, and Sama Formations and about two-thirds (700 m) of the Iscayachi Formation. HCS—hummocky cross-stratification; isp.—ichnospecies.
Life and death of a Cambrian–Ordovician basin: An Andean three-act play
515
TABLE 1. ORDOVICIAN AND CAMBRIAN STRATIGRAPHY OF NORTHWESTERN ARGENTINA AND SOUTHERN BOLIVIA Period Stratigraphy of Stratigraphy of northwestern Argentina southern Bolivia Ordovician
Santa Rosita Formation
Iscayachi Formation
Cambrian
Calhualmayoc Formation Campanario Formation Lizoite Formation
Sama Formation Torohuayco Formation Camacho Formation
WEST
EAST
Middle Ordovician
Challa Mayu 125 m
Isograptus victoriae
Azygograptus lapworthi
0m
Chaupiuno 1+2 500 m
Pilar Punta Loma
1000 m
Azygograptus lapworthi
DELTA
1000 m 0m
Rio Palomita
Mal Paso
Sella 1
(Pirancha, Sella, and Rumi Orkho Formations) 500 m
Baltograptus minutus
500 m
500 m
upper Lower Ordovician
Cieneguillas 2 Baltograptus minutus
500 m
0m
1000 m
Cieneguillas 1
0m
(Obispo and Taraya Formations)
Expansograptus protobalticus
Abra Negra
Expansograptus protobalticus
Expansograptus holmi
Taraya
0m
turbidite ramp Expansograptus holmi
0m
2000 m
Tetragraptus phyllograptoides
1000 m
Culpina 2 0m
500 m
Hunnegraptus copiosus
Tetragraptus phyllograptoides
Araneograptus murrayi
H. copiosus
A. murrayi
Tremadoc
0m
Aorograptus victoriae + Adelograptus sp.
1000 m
0m
Incahuasi Culpina 1
Aorograptus victoriae + Adelograptus sp.
Sama Rhabdinopora flabelliformis
Key
1000 m Lower limit graptolite zone certain
Rhabdinopora flabelliformis
0m
Lower limit graptolite zone uncertain Mud- and siltstones
no graptolites Sandstones Reverse fault
DELTA
no graptolites
(Iscayachi Formation) 0m
Figure 3. Transect through the Lower Ordovician shelf in southern Bolivia showing the twelve investigated sections (for location see Fig. 1) and their biostratigraphic correlation based on graptolites (Egenhoff et al., 2004). Note the enormous thickness of both the turbidite ramp and the delta complex succession. Fm.—Formation.
hiatus
Figure 4. Stratigraphic scheme for the Lower Ordovician succession of southern Bolivia on an east-west transect between Tarija or Sella and Tupiza (Fig. 1). Note the diachronous relationships of formations. Fm.—Formation.
Tremadoc
Obispo Fm. Culpina Incahuasi Sama
Taraya
Obispo Fm.
Taraya Fm.
Abra Negra
Sella & Rumi Orkho Fms.
Sella
m.
Mal Paso
Pilar Punta Loma
Pircancha, Sella, and Rumi Orkho Formations Obispo (O) and Taraya (T) Formations
Middle Ordovician
Iscayachi Fm.
Lower Ordovician
Rio Palomita
cha F
Chaupiuno
Pircan
Ordovician
east
Challa Mayu
Cieneguillas
west
Middle Ordovician
Egenhoff
Upper Lower Ordovician
516
1800 m (Rivas et al., 1969)
Silty pelites with abundant bioturbation that gradually grade into sandstone-rich facies toward the top of the formation; HCS, SCS, wave-, and combinedflow ripples are common in the upper part; lenticular sandstone bodies are eroded into peltic and sand-rich facies near the top
Prodelta pelitic to sandy delta front deposits showing wave influence, with some delta channels near the top Prograding wavedominated delta
T O
O Obispo Fm., 2080 m minimum, thicking toward the W (Egenhoff, 2000) T Taraya Fm., 490 m (Egenhoff, 2000)
Obispo Fm Black siliciclastic mudstones with mm- to cm-thick siltstone Deposition from laminae and dm-thick sandbackground stone intercalations; sandstones sedimenation, and graded with structureless lower high- to low-density part, faint laminated central and turbidity currents on current ripple cross-bedded top; turbidite ramp individual sandstone beds are laterally traceable for several hundreds of meters; pelites generally lack bioturbation Taraya Fm Black siliciclastic mudstones Sedimentation from with or without sub-mm-thick suspension, siltstone laminae, no sandstone deposition on intercalations elevated horst block on turbidite ramp
Figure 5. Generalized section with facies descriptions and sedimentological interpretations through the Lower Ordovician succession in southern Bolivia comprising the Obispo, Taraya, Pircancha, Sella, and Rumi Orkho Formations. The Obispo Formation conformably overlies the Iscayachi Formation (Fig. 2). Fm.—Formation.
Life and death of a Cambrian–Ordovician basin: An Andean three-act play
A
B Abra Negra
Cieneguillas
A
N
B
A´
Sama
Tupiza
Taraya
Chaupinuo Cieneguillas Sella B´
Tupiza
Tarija
Tarija
Taraya horst block
N
Abra Tupiza Negra
Taraya
A
B
Chaupinuo Cieneguillas Sella A´´ B´´ Pilar Punta Tupiza Tarija Mal Loma position of Paso coastline during
BOLIVIA
ARGENTINA
ARGENTINA
50 km
50 km
50 km
Cieneguillas Sama
Rio Palomita Challa Mayu
N
BOLIVIA
BOLIVIA ARGENTINA
C
517
A´ Tupiza
A´
B´ Chaupiuno
Azygograptus lapworthi Biozone
Mal Tupiza Paso
Sella
B´ Chaupiuno
Sella
delta
delta
turbidite ramp
D
Key turbidite ramp prodelta facies land
Locality Political border
Challa Mayu A´´´
B´´´
Tupiza N
Figure 6. Evolution of the shelf in southern Bolivia in the Early Ordovician, shown in four steps in schematic maps with corresponding cross-sections. (A) During the Middle and Upper Tremadocian and parts of the late Early Ordovician, turbidite ramp deposition dominated in the study area. A horst block with condensed pelite sedimentation characterized the locality Taraya and continued farther to the south (Egenhoff, 2000). (B) In the Expansograptus holmi graptolite biozone (late Early Ordovician), prodelta sedimentation occurred at Cieneguillas reflecting the onset of delta progradation. (C) During the Azygograptus lapworthi graptolite biozone (early Middle Ordovician), delta front facies and channels were exposed at Chaupiuno, reflecting the position of the coastline in that area. (D) In the Isograptus victoriae graptolite biozone (early Middle Ordovician), prodelta sediments were exposed in Challa Mayu near Tupiza, indicating that the delta prograded ~100 km to the west.
approximate position of coastline during Isograptus victoriae Biozone
Tarija BOLIVIA
50 km
ARGENTINA
A´´´
B´´´
Mal Tupiza Paso delta
518
Egenhoff
running north-south and interfingers with the Obispo Formation toward the east and west, as reflected in successive intercalation of turbidite sandstones and a dramatic increase in the thickness of individual graptolite biozones. The Taraya Formation represents deposition on an elevated horst block rather than a distal condensed basin facies. The evidence is that this unit grades laterally into the Obispo Formation toward the proximal and distal areas and that even west of the horst block sediment transport was directed toward the west, thereby ruling out an origin of these turbidites from the western basin margin. Furthermore, this northsouth-trending tectonic block is still reflected in the sediment distribution patterns of the overlying Pircancha Formation (upper Lower Ordovician), where individual turbidites originating from the east are deflected to the north by the horst structure. Throughout the Early Ordovician, no prodelta sediments were intercalated into the Obispo ramp succession, indicating that the system was purely aggradational despite an extremely high turbidite sediment input. The accommodation space created by sediment loading combined with tectonic subsidence prevented progradation of the adjacent delta over the Obispo ramp until the late Early Ordovician (Fig. 6). But by the time of deposition of the Expansograptus holmi graptolite biozone (later in late Early Ordovician time), the delta started to build out toward the west and successively filled up the accommodation space. By the early Middle Ordovician, the delta in southern Bolivia had advanced ~100 km to the west and was located a few tens of kilometers east of Tupiza (Figs. 1 and 6), and much of the shelf area submerged during the Lower Ordovician was exposed. The diachronous formation boundary between the turbidite ramp Obispo Formation and the overlying deltaic Pircancha, Sella, and Rumi Orkho Formations illustrates this progradational delta pattern (Fig. 4). Individual upper Lower Ordovician and lower Middle Ordovician graptolite zones are generally thicker during the progradational phase in Obispo ramp sediments (localities Cieneguillas, Chaupiuno, Mal Paso, Pilar Punta Loma, and Rio Palomita) in comparison to the preceding aggradational phase (localities Cieneguillas and Culpina; Figs. 1 and 3). For example, the Baltograptus minutus biozone is 450 m thick in the east near Sella, but ~1600 m thick at Cieneguillas and Chaupiuno farther to the west (Figs. 1 and 3). The same probably applies to the Azygograptus lapworthi biozone, but cannot be definitely proven because the boundary with the overlying Isograptus victoriae biozone is not preserved at any locality. These differences in thickness show that subsidence decreased on the shelf during the late Early Ordovician successively from east to west as a consequence of (1) the reduction of accommodation space in the proximal zone so that more sediment was shed onto the deep shelf, (2) an increase in sediment input from the hinterland, or (3) a combination of the two. Accumulation of greater thicknesses of sediment on the deep shelf during the progradational phase should have led to an increase in sediment load–induced subsidence. However, as the accommodation space was successively reduced, it is likely that tectonically induced uplift accounted for the shallowing of facies throughout the Pircancha Formation as well as the thinning of
individual graptolite biozones eastward within the Lower to Middle Ordovician shelf succession. This profound change affected the entire eastern shelf during the late Early and early Middle Ordovician; the westward shift of the coastline is observed in both northwestern Argentina (Bahlburg, 1990) and Bolivia. In the late Early Ordovician, the western part of the basin adjacent to the eastern shelf deepened substantially and an elongated north-south-trending foreland trough formed (Egenhoff, 2000). Deep marine foreland basin strata of this early subsidence stage—the Didymograptus bifidus graptolite biozone (Bahlburg et al., 1990), which corresponds to the Baltograptus minutus graptolite biozone of Bolivia (J. Maletz, personal commun., 2005)—crop out exclusively in northwestern Argentina (Bahlburg, 1991) and are covered by younger foreland sediments in southern Bolivia. But as basin evolution was triggered by the movements of the same terrane delimiting this trough in the west (Bahlburg and Hervé, 1997), the formation of the foreland basin started, probably simultaneously in northwestern Argentina and southern Bolivia. It is likely that the initiation of the foreland basin and the onset of delta progradation on the shelf started at the same time and are therefore related to the same tectonic event in the late Early Ordovician. This scenario is well explained with models of foreland basin evolution that succeeded a passive margin or rift stage (Cohen, 1982; Stockmal et al., 1986). With the onset of compressional tectonic forces, the flexure of the overridden plate produced a deepening of the basin adjacent to the load responsible for the flexure—the deep marine foreland trough exposed in northwestern Argentina (Bahlburg, 1990). Simultaneously, a forebulge that developed farther inland on the overridden plate accounts for the reduction of accommodation space on the Early Ordovician shelf (cf. Jacobi, 1981). A similar scenario has been proposed for the Ordovician basin evolution in Argentina (Bahlburg, 1990, 1991). However, Bahlburg (1990, 1991) proposed a reversal from extension to compression during the Darriwilian (in his texts these were still late Arenig or around the ArenigLlanvirn boundary), but at that time the foreland basin itself as well as the forebulge was already fully evolved. BASIN CLOSURE The Upper Ordovician succession has been studied in the localities of Mauka Huasi, Rio Marquina, San Pablo de Lipez, Soniquera, and Tica Tica (Fig. 1). Furthermore, the interpretations presented rely on the thickness distribution of Hirnantian strata (Fig. 7; Suarez Soruco, 1995). A summary of the facies and characteristics of the Upper Ordovician succession is given in Figure 8. In the early Late Ordovician, the southern Bolivian foreland basin was the site of turbidite deposition (Egenhoff, 2000). The stratigraphic succession coarsens from east to west (Fig. 9). At Mauka Huasi (Fig. 1) it is dominated by pelitic strata; at Rio Marquina, sand-rich turbidites are common, and flute marks are oriented to the north (Fig. 9). The Altiplano outcrops at San Pablo de Lipez and Soniquera (Fig. 1) are composed of sandy to conglomeratic turbidite beds with sedimentary structures
Life and death of a Cambrian–Ordovician basin: An Andean three-act play
MOVEMENTS OF THE AREQUIPA-ANTOFALLA TERRANE RELATIVE TO GONDWANA The sedimentary filling of the Cambrian–Ordovician basin in northwestern Argentina and southern Bolivia reflects the movements of the Arequipa-Antofalla terrane relative to Gondwana. Extensional tectonics between the terrane and Gondwana started in the Cambrian in the north and spread south (cf. Gohrbandt, 1992). From the Cambrian to the Early Ordovician, the basin widened more in Bolivia than in Argentina, and although no ocean crust formed (Bahlburg et al., 1997), the “mainland” of the Arequipa-Antofalla terrane was separated from Gondwana by a wider sea in Bolivia than in Argentina. This V-shaped basin
68°
64°
66°
16°
BOLIVIA 0m 0m
0m
18°
0m
22°
200m
more than 1400 m in Tica Tica
20°
10 0 60 0m 0m
20
CHILE
indicating sediment transport toward the east. This facies differentiation reflects the foreland basin morphology, with a steep western margin from which sand-rich turbidites were shed downslope toward the east (Fig. 9). In the basin center, represented by the Rio Marquina locality, the orientation of sedimentary structures indicates that the turbidity currents were redirected to the north, parallel to the basin axis. Toward the end of the Ordovician, the basin underwent a pronounced shallowing, and the Lower and middle Upper Ordovician sediments, represented by the Marquina, Angosto, Kollpani, and Tapial Formations (Fig. 8; Müller et al., 1996), were covered by a Hirnantian succession up to 1400 m thick that entirely filled the available accommodation space. The bulk of the rocks are the diamictites of the Cancañiri Formation (Fig. 8; Sempere, 1995). These sediments originated from inland glaciers that record the Hirnatian ice age on Gondwana (Sempere, 1995; Suarez Soruco, 1995). However, it remains a matter of debate whether all of the diamictites observed in southern Bolivia are true tillites or have been at least partly remobilized (Diaz Martínez, 1997; Schönian, 2000). Massive sandstones intercalated with pebbly mudstones with microfossils (Fig. 8) in the Tica Tica locality (Figs. 1 and 7) confirm the reworking and input of nonglacial material into this basin. The thickness distribution of the Hirnantian Cancañiri Formation reflects the asymmetric foreland basin geometry, with steep western and gentle eastern margins (Fig. 7). In a north to south direction, the thickness distribution indicates that the basin was deepest in the locality of Tica Tica, the area that experienced the greatest subsidence. Toward the south, the basin center gradually shallowed, which is reflected in successively thinner strata, with the Cancañiri Formation pinching out a few kilometers north of the Bolivian-Argentinean border. In the Puna of northwestern Argentina, folded lower Upper Ordovician strata are unconformably overlain by Silurian sediments (Benedetto and Sanchez, 1990). A compressional tectonic event (the “Ocloyic orogeny”) responsible for the folding of Ordovician strata had therefore occurred prior to the deposition of the Silurian sediments, which is best explained by a closure of the basin in northwestern Argentina. In Bolivia, in contrast, the basin remained open until the end of the Ordovician, as reflected in the deposition of the Hirnantian strata.
519
TUPIZA
ARGENTINA Figure 7. Thickness distribution of the Ashgill Cancañiri Formation in Bolivia, based on Suarez Soruco (1995). In southern Bolivia, it reflects the latest Ordovician basin geometry with a steep western and a lowinclined eastern basin flank. The thinning of Hirnantian strata to the south mirrors increased tectonic compression toward that direction.
geometry indicates that the terrane rotated counterclockwise relative to Gondwana, with a rotation pole located in northwestern Argentina during the extensional phase (Fig. 10). These reconstructions based on sedimentological and facies evidence are at odds with those based on paleomagnetic data used by Forsythe et al. (1993), who suggested a rotation pole in southern Peru for the initial movement of the terrane. The switch from an extensional to a compressional regime in the late Early Ordovician (the Expansograptus holmi biozone) defines the moment when the Arequipa-Antofalla terrane reversed direction and again approached Gondwana. When compressional tectonics started, the basin already had an inherited V-shaped geometry and was narrower in the south than in the north. The collision between the terrane and Gondwana was therefore significantly greater in northwestern Argentina than in southern Bolivia, resulting in the Ordovician deformation of
Hirnantian
Canacañiri Formation Marquina, Angosto, Kollpani, and Tapial Formations
Egenhoff
Lower and middle Upper Ordovician
520
0 - 1400 m (Suarez Soruco, 1995)
> 5500 m (Müller et al., 1996)
Diamictites with some massive sandstone and pebbly mudstone intercalations
Deposition from glaciers; enormous thickness of more than1000 m argues for intercalated debris flows consisting of reworked glacial material; pebbly mudstones also indicate debris flow deposition; massive sandstones are the product of rapid deposition, likely in a standing water body; late basin fill, marine (?) and terrestrial facies
Sandstone-mudstone intercalations, succession with Deposition in turbidite a distinct fining trend; basin; lenticular sandm u d s t o n e s l a m i n a t e d , stone bodies represent no bioturbation; sand- channels, sheet-like stone graded, individual sandbodies are lobes; beds traceable for kilo- slumps represent areas meters; in the lower portion, of locally high insandstone units with lenti- clination; mudstones cular geometry; in the represent background central part, more sheet-like sedimentation and mud geometries; some slumps turbidites; early foreunits, tens of meters thick, land basin underfilled stage intercalated into the central part
Figure 8. Overview section of the Upper Ordovician succession in southern Bolivia with facies descriptions and sedimentological interpretations, based on Suarez Soruco (1995), Müller et al. (1996), and the author’s own data. The lower part of the Upper Ordovician turbidite succession is overlain by Hirnantian terrestrial to shallow marine sediments, indicating complete filling of the accommodation space.
the “Ocloyic orogeny” (Turner, 1960; Coira et al., 1982; Mon and Hongn, 1991). During the compressional phase, the terrane rotated clockwise with respect to Gondwana. However, because the basin closed completely in the south but remained open in the north, it is likely that the rotation pole was located in southern Peru during the evolution of the foreland basin, as is also suggested by paleomagnetic data (Fig. 10; Forsythe et al., 1993). Tectonic movements within the basin between the ArequipaAntofalla terrane and Gondwana likely continued at least until the end of the Ordovician. This is indicated by the locally very thick Hirnantian succession, for example, in Tica Tica. The form of the Silurian successor basin, with a connection to the open ocean through northern Bolivia and Peru similar to that of its precursor (Gohrbandt, 1992), suggests that subsidence in great parts of the Cambrian–Ordovician basin continued beyond the Hirnantian.
DISCUSSION Timing of Basin Initiation The timing of foreland basin initiation is determined from graptolite biostratigraphy. The onset of delta progradation on the cratonward basin margin interpreted to reflect initiation of tectonic compression occurred in the Expansograptus holmi biozone (Egenhoff et al., 2004). The earliest sediments identified as foreland basin deposits are about half a graptolite zone older (Didymograptus bifidus biozone; Bahlburg et al., 1990), a difference in timing that corresponds to ~1–1.5 m.y. according to Cooper and Lindholm (1990) and Egenhoff (2000). This difference in relative ages of the oldest sediments recognized as foreland basin deposits on the cratonward shelf in contrast to
Life and death of a Cambrian–Ordovician basin: An Andean three-act play
W
521
E western slope
Arequipa-Antofalla Block
basin center
eastern slope
(sandstone and (sandstoneconglomerate rich facies) facies, high- inclined margin)
(mudstone-dominated facies, low-inclined margin)
Mauka Huasi
Soniquera
Tupiza
San Pablo de Lipez
Figure 9. Schematic reconstruction of the foreland basin in southern Bolivia in the early Late Ordovician. The Soniquera and San Pablo de Lipez localities (Fig. 1) represent the coarse-grained western basin margin, where sediment is shed toward the east into the basin center. Rio Marquina reflects northward sediment transport in the deepest part parallel to the basin axis. Mauka Huasi represents the fine-grained and low-inclined eastern flank. (A) Ripple mark values from Soniquera; (B) flute cast measurements from San Pablo de Lipez; (C) flute cast measurements from Rio Marquina.
Rio Marquina
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Figure 10. Schematic reconstruction of the movements of the Arequipa-Antofalla terrane (AAT) during the Cambrian–Ordovician, based on basin fill geometries. (A) During the initial extensional period, the basin opened successively from north to south, mirroring the counterclockwise rotation of the terrane relative to Gondwana with an Euler pole in northwestern Argentina. (B) From the late Early Ordovician (Expansograptus holmi graptolite biozone) on, the basin started to narrow when the terrane approached Gondwana again. During the Hirnantian, the basin closed in the south, resulting in the “Ocloyic orogeny,” but in southern Bolivia the basin remained open. The Euler pole for the clockwise movement of the terrane from the late Early Ordovician to the Hirnantian was located in Peru (Forsythe et al., 1993).
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the deep marine basin is best explained by ecological factors. Graptolites occur much more frequently in the Lower Ordovician shelf succession than in the deep marine foreland basin strata, which is probably a function of greater sediment input, thereby producing a higher dilution of graptolites with sediment. The onset of compressional tectonics is therefore more easily and precisely determined on the gently inclined cratonward shelf due to a better biostratigraphic control. Compression is indicated there by the onset of a long-term reduction of accommodation space that reflects the formation of a forebulge (for adjacent northwestern Argentina, see Bahlburg 1990, 1991; Astini, 2003) and an overall narrowing of the entire basin (cf. Allen and Allen, 2005). With a highstand of second order in the Early Ordovician (Ross and Ross, 1992; Nielsen, 2003), global sea-level changes cannot account for accommodation reduction in southern Bolivia and northwestern Argentina. Timing of Reversal of the Basin-Tectonic Regime The crucial argument for a change in direction of ArequipaAntofalla terrane movement is the switch of the geotectonic regime in the Cambrian–Ordovician basin in Argentina and Bolivia. In the present article, this transition from an extensional setting to a foreland basin is proposed to have occurred in the late Early Ordovician. Conversely, Bahlburg (1990, 1991) argued for a switch in basin type in the Darriwilian. He used the following evidence to support his interpretation: (1) decrease of the content of volcaniclastic material within the turbidite input into the basin from the Early Ordovician to the Darriwilian, which he interprets to be the result of the extinction of an arc due to the onset of compression; (2) rapid subsidence estimates in the entire basin during the studied interval (late Early Ordovician to early Late Ordovician), which is more indicative of foreland basins; and (3) an upward-fining and -deepening succession observed in the eastern Puna, which he interpreted as resulting from a reversal of the geotectonic regime. The dominance of volcanic-derived debris in the upper Lower Ordovician Puna sediments undoubtedly argues for the erosion of a magmatic arc as proposed by Bahlburg (1990, 1991). However, it remains speculative whether the arc was still active during the late Early Ordovician, which represents the age of the oldest volcaniclastics found so far in this basin. According to Breitkreuz et al. (1989) and in contrast to Bahlburg (1990, 1991), the composition of the upper Lower Ordovician portion of the Puna Turbidite Complex (sensu Bahlburg, 1990) indicates a rapid decrease of input of volcaniclastic debris into the basin, which would rather argue for the successive erosion of a remnant arc. Therefore, the exact timing of arc extinction remains unclear and cannot be deduced from the lithology of the basin fill. It is therefore more likely that the arc became inactive in the late Early Ordovician, which would parallel a general switch of magmatic activity in the northern Puna (Coira et al., 1999). Bahlburg’s (1990, 1991) subsidence curve from northwestern Argentina shows the highest rates in the upper Lower and
lower Middle Ordovician, with a subsequent decrease in the Darriwilian and the Upper Ordovician. This is in good agreement with onset of a rapidly subsiding foreland basin in the late Early Ordovician (this article). The upward-fining trend observed by Bahlburg (1990, 1991) was very likely produced by the successively deepening foreland basin, as is also reflected in the high initial subsidence values. However, it does not contradict a possible earlier formation of the foreland basin with the change of geotectonic regimes in the late Early Ordovician. All available data indicate that the northwestern Argentinean part of the basin was in a back-arc position during the extensional phase (Bahlburg, 1990, 1991; Moya et al., 1993; Zimmermann and Bahlburg, 2003). In southern Bolivia, only the eastern flank of this extensional basin is exposed; there sediment was derived exclusively from the Brazilian Shield in the east, consisting of reworked Proterozoic and older rocks. The lack of volcanicderived debris in the basin fill can therefore not be taken as an argument to prove the absence of an arc in southern Bolivia, and it is not possible to prove whether the Lower to Middle Ordovician volcanic arc of northwestern Argentina continued into southern Bolivia. So although no arc signature is reflected in the geochemistry of the Ordovician sediments (Egenhoff and Lucassen, 2003), the Bolivian extensional basin is regarded as a backarc basin based on the studies from northwestern Argentina. CONCLUSIONS The outcrop study of the large-scale sedimentary architecture of the Bolivian Cambrian–Ordovician succession enabled recognition and reconstruction of terrestrial to shallow marine, deltaic, deep marine turbidite, and glacial depositional environments in this Andean area. The juxtaposition of these environments and stratigraphic relationships reflects the entire history of a basin that developed following rifting and successive reamalgamation of the Arequipa-Antofalla terrane from Gondwana. The basinal setting with differential waxing and waning of depocenters is further corroborated by data from northwestern Argentina (Moya, 1988; Bahlburg, 1990, 1991). The evolution of the basin can be divided into three stages: (1) initial extensional movements between the Arequipa-Antofalla terrane and Gondwana, leading to the formation of an elongated depositional trough (the successive opening of the basin from north to south during the Cambrian reflects the initial counterclockwise rotation of the terrane, with a rotation pole located in northwestern Argentina); (2) transition of an extensional to a foreland basin with ensuing reversal of the tectonic regime to compressional dynamics; and (3) reaccretion of the terrane onto Gondwana, resulting in partial basin closure (the south to north progressive basin closure during the Hirnantian indicates a clockwise rotation of the terrane relative to Gondwana, with a Euler pole in southern Peru). The thinning of Cambrian to Lower Ordovician strata to the south reflects decreasing subsidence within this basin and serves
Life and death of a Cambrian–Ordovician basin: An Andean three-act play as evidence of stronger extensional tectonic forces in the northern part of the basin in southern Bolivia compared to northwestern Argentina. This differential tectonic regime produced a V-shape Cambrian–Ordovician basin that widened toward the north. The transition from an extensional to a compressional basin occurred in the late Early Ordovician (Expansograptus holmi biozone). Foreland trough formation is reflected in the profound deepening of the western side of the basin, the onset of delta progradation on the eastern shelf as a consequence of reduced accommodation and forebulge formation, and the resulting significant narrowing of the basin. Most reliable biostratigraphic data from the interval of tectonic reversal comes from the distal cratonward margin. Planktic graptolites display a higher diversity and better stratigraphic resolution and occur in far greater abundance on the shelf than in the adjacent basin (cf. Finney and Berry, 1997). Graptolite biostratigraphy from shelf facies allows the change in basin geometry to be correlated with the precision of one graptolite biozone, an accuracy that would not have been possible using the faunal content of the basinal succession alone. The Late Ordovician foreland basin fill records a shallowing of the environment from deep marine turbidite sedimentation to the deposition of partially terrestrial glacial deposits. The geometry of the Hirnantian sediments reflects the closure of the basin in the south in northwestern Argentina, while in southern Bolivia the basin remained open and accumulated more than 1000 m of sediments. ACKNOWLEDGMENTS My research was made possible by funding from the German Research Foundation DFG, projects SFB 267, TP C3, and EG 141/1-1. Many thanks to Jörg Maletz, Buffalo, for inspiring discussions, and to Arndt Peterhänsel, Potsdam, and Sally Sutton, Fort Collins, for correcting the English. Special thanks go to Mario Salamanca, Sucre, for help in the field, and to Susanne Hesse, Freiberg, for drawing several of the figures. Luis Buatois, Saskatoon, and Stanley Finney, Long Beach, provided very helpful and detailed reviews that significantly improved the quality of the manuscript. REFERENCES CITED Ahlfeld, F., and Branisa, L., 1960, Geología de Bolivia: La Paz, Bolivia, Editorial Don Bosco, 331 p. Allen, P.A., and Allen, J.R., 2005, Basin analysis: Principles and applications: Oxford, Blackwell, 449 p. Araníbar, O., 1979, Estudio geológico de la parte sur de la hoja Padcaya N°6629.: Servicio Geológico de Bolivia (internal report of the Geological Survey, no pagination). Astini, R.A., 2003, The Ordovician proto-Andean basins, in Benedetto, J.L., ed., Ordovician fossils of Argentina: Secretaria de Ciencia y Tecnología, Universidad Nacional de Cordoba, p. 1–74. Bahlburg, H., 1990, The Orovician basin in the Puna of NW Argentina and N Chile: Geodynamic evolution from back-arc to foreland basin: Geotektonische Forschungen, v. 75, p. 1–107. Bahlburg, H., 1991, The Ordovician back-arc to foreland successor basin in the Argentinian-Chilean Puna: Tectono-sedimentary trends and sea-level
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changes, in MacDonald, D.I.M., ed., International Association of Sedimentologists (Oxford) Special Publication 12, p. 465–484. Bahlburg, H., and Hervé, F., 1997, Geodynamic evolution and tectonostratigraphic terranes of northwestern Argentina and northern Chile: Geological Society of America Bulletin, v. 109, p. 869–884, doi: 10.1130/00167606(1997)109<0869:GEATTO>2.3.CO;2. Bahlburg, H., Breitkreuz, C., Maletz, J., Moya, M.C., and Salfity, J.A., 1990, The Ordovician sedimentary rocks in the northern Puna of Argentina and Chile: New stratigraphic data based on graptolites: Newsletters on Stratigraphy, v. 23, p. 69–89. Bahlburg, H., Kay, S.M., and Zimmermann, U., 1997, New geochemical and sedimentological data on the evolution of the Early Paleozoic Gondwana margin in the southern Central Andes, in Geological Society of America Annual Meeting (Salt Lake City, Utah, 20–23 October), v. 29, no. 6, p. 378. Benedetto, J.L., and Sanchez, T.M., 1990, Fauna y edad del estratotipo Salar del Rincón (Eopaleozoico, Puna Argentina): Ameghiniana, v. 27, p. 317–326. Breitkreuz, C., Bahlburg, H., Delakowitz, B., and Pichowiak, S., 1989, Paleozoic volcanic events in the Central Andes: Journal of South American Earth Sciences, v. 2, p. 171–189, doi: 10.1016/0895-9811(89)90045-X. Buatois, L.A., and Mángano, M.G., 2003, Sedimentary facies, depositional evolution of the upper Cambrian–lower Ordovician Santa Rosita Formation in northwest Argentina: Journal of South American Earth Sciences, v. 16, no. 5, p. 343–363, doi: 10.1016/S0895-9811(03)00097-X. Cohen, C.R., 1982, Model for a passive to active continental margin transition: Implications for hydrocarbon exploration: American Association of Petroleum Geologists Bulletin, v. 66, p. 708–718. Coira, B., Davidson, J., Mpodozis, C., and Ramos, V., 1982, Tectonic and magmatic evolution of the Andes of northern Argentina and Chile: Earth-Science Reviews, v. 18, p. 303–332, doi: 10.1016/0012-8252(82)90042-3. Coira, B., Pérez, B., Flores, P., Kay, S.M., Woll, B., and Hanning, M., 1999, Magmatic sources and tectonic setting of Gondwanan margin Ordovician magmas, northern Puna of Argentina and Chile, in Ramos, V.A., and Keppie, J.D., eds., Laurentia-Gondwana connections before Pangea: Boulder, Colorado, Geological Society of America, Special Paper 336, p. 145–170. Cooper, R.A., and Lindholm, K., 1990, A precise worldwide correlation of early Ordovician graptolite sequences: Geological Magazine, v. 127, p. 497–525. Diaz Martínez, E., 1997, Facies y ambientes sedimentarios de la formacion Cancañiri (Silurico inferior) en La Cumbre de La Paz, norte de la Cordillera Oriental de Bolivia: Geogaceta, v. 22, p. 55–57. Egenhoff, S., 2000, Sedimentologie und Beckenentwicklung im Ordovizium in Südbolivien: Berlin, Berliner Geowissenschaftliche Abhandlungen, 173 p. Egenhoff, S.O., and Lucassen, F., 2003, Chemical and isotope composition of Lower to Upper Ordovician sedimentary rocks (Central Andes / south Bolivia): Implications for their source: Journal of Geology, v. 111, p. 487– 497, doi: 10.1086/375280. Egenhoff, S., Maletz, J., and Erdtmann, B.-D., 2004, Lower Ordovician graptolite biozonation and lithofacies of southern Bolivia: Relevance for palaeogeographic interpretations: Geological Magazine, v. 141, no. 3, p. 287–299, doi: 10.1017/S0016756804009239. Finney, S., 2005, Global series and stages for the Ordovician system: A progress report: Geologica Acta, v. 3, p. 309–316. Finney, S.C., and Berry, W.B.N., 1997, New perspectives on graptolite distribution and their use as indicators of platform margin dynamics: Geology, v. 25, p. 919–922. Forsythe, R.D., Davidson, J., Mpodozis, C., and Jesinkey, C., 1993, Lower Paleozoic relative motion of the Arequipa Block and Gondwana: Paleomagnetic evidence from the Sierra de Almeida of northern Chile: Tectonics, v. 12, p. 219–236. Gohrbandt, K.H.A., 1992, Paleozoic paleogeographic and depositional developments on the central proto-Pacific margin of Gondwana: Their importance to hydrocarbon accumulation: Journal of South American Earth Sciences, v. 6, no. 4, p. 267–287, doi: 10.1016/0895-9811(92)90046-2. Jacobi, R.D., 1981, Peripheral bulge—A causal mechanism for the Lower Ordovician unconformity along the western margin of the northern Appalachians: Earth and Planetary Science Letters, v. 56, p. 245–251, doi: 10.1016/0012-821X(81)90131-X. Kley, J., 1993, Der Übergang vom Subandin zur Ostkordillere in Südbolivian (21°15–22°S): Geologische Struktur und Kinematik [doctoral thesis]: Berlin, Freie Universität, 88 p.
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Kumpa, M., and Sanchez, M.C., 1988, Geology and sedimentology of the Cambrian Grupo Meson (NW Argentina), in Bahlburg, H., Breitkreuz, C., and Giese, P., eds., The southern Central Andes: Lecture Notes in Earth Sciences: Berlin, Springer, p. 39–53. Loewy, S.L., Conelly, J.N., and Dalziel, I.W.D., 2004, An orphaned basement block: The Arequipa-Antofalla Basement of the central Andean margin of South America: Geological Society of America Bulletin, v. 116, no. 1/2, p. 171–187, doi: 10.1130/B25226.1. Mon, R., and Hongn, F., 1991, The structure of the Precambrian and Lower Paleozoic Basement of the Central Andes between 22° and 32° LAT: Geologische Rundschau, v. 80, no. 3, p. 745–758, doi: 10.1007/BF01803699. Moya, M.C., 1988, Lower Ordovician in the southern part of the Argentine Eastern Cordillera, in Bahlburg, H., Breitkreuz, C., and Giese, P., eds., The southern Central Andes: Lecture Notes in Earth Sciences: Berlin, Springer, p. 55–69. Moya, M.C., 2002, The Ordovician basin of northern Argentina, in Aceñolaza, F.G., ed., Aspects of the Ordovician System in Argentina: INSUGEO, Serie Correlación Geológica, v. 16, p. 281–294. Moya, M.C., Malanca, S., Hongn, F.D., and Bahlburg, H., 1993, El Tremadociano temprano en la Puna occidental argentina: Decimosegundo Congreso Geológico Argentino y segundo Congreso de Hidrocarburos Actas, v. 2, p. 20–30. Moya, M.C., Malanca, S., Monteros, J.A., and Cuerda, A., 1994, Bioestratigrafía del Ordovicico Inferior en la Cordillera Oriental Argentina basada en graptolitos: Revista Espanola de Paleontologia, v. 9, p. 91–104. Müller, J.P.M., Maletz, J., Egenhoff, S., and Erdtmann, B.D., 1996, Turbiditas Caradocianas– (?)Ashgillianas inferiores en la Cordillera Oriental al sur de Bolivia: Implicaciónes cinemáticas, in Suarez Soruco, R., ed., XII Congreso Geológico de Bolivia, Tarija, p. 747–753. Nielsen, A.T., 2003, Ordovician sea-level changes: Potential for global event stratigraphy, in International Symposium on the Ordovician System, San Juan, Argentina, p. 445–449. Pankhurst, R.J., and Rapela, C.W., 1998, The proto-Andean margin of Gondwana: Geological Society of London Special Publication 142, 383 p.
Reutter, K.-J., Scheuber, E., and Wigger, P.J., 1994, Tectonics of the southern Central Andes: Berlin, Springer, 333 p. Rivas, S., Fernandez, C., and Alvarez, R., 1969, Estratigrafia de los sistemas Ordovicico–Cambrico y Precambrico en Tarija, Sud de Bolivia: Sociedad Geologica Boliviana Boletín, v. 9, p. 27–50. Rodrigo, L.A., and Castaños, A., 1978, Sinopsis estratigráfica de Bolivia, parte 1: La Paz, Academia Nacional de Ciencias, 116 p. Ross, J.R.P., and Ross, C.A., 1992, Ordovician sea-level fluctuations, in Webby, B.D., and Laurie, J.R., eds., Global perspectives on Ordovician Geology: Rotterdam, Balkema, p. 327–335. Schönian, F., 2000, Die Ablagerungsgeschichte der Cancañiri-Diamiktite (Ordovizium/Silur) im Gebiet der Santa Cruz Loma bei Tarija, Südbolivien [diploma thesis]: Berlin, Humboldt Universität, 129 p. Sempere, T., 1995, Phanerozoic evolution of Bolivia and adjacent regions, in Tankard, A.J., Soruco, S., and Welsink, H. J., eds., Petroleum basins of South America: American Association of Petroleum Geologists Memoir 62, p. 207–230. Stockmal, G.S., Beuamont, C., and Boutilier, R., 1986, Geodynamic models of convergent margin tectonics: Transition from rifted margin to overthrust belt and consequences for foreland-basin development: American Association of Petroleum Geologists Bulletin, v. 70, p. 181–190. Suarez Soruco, R., 1995, Comentarios sobre la edad de la Formacion Cancañiri: Revista Tecnica de Yacimientos Petroliferos Fiscales Bolivianos, v. 16, p. 51–54. Turner, J.C.M., 1960, Estratigrafía de la Sierra de Santa Victoria y adjacencias: Academia Nacional de Ciencias de Córdoba Boletín, v. 41, p. 163–196. Zimmermann, U., and Bahlburg, H., 2003, Provenance analysis and tectonic setting of the Ordovician clastic deposits in the southern Puna Basin, NW Argentina: Sedimentology, v. 50, p. 1079–1104, doi: 10.1046/j.13653091.2003.00595.x.
MANUSCRIPT ACCEPTED BY THE SOCIETY 3 OCTOBER 2006
Printed in the USA
Geological Society of America Special Paper 423 2007
A Late Ordovician ice sheet in South America: Evidence from the Cancañiri tillites, southern Bolivia Frank Schönian* Museum of Natural History, Humboldt–University of Berlin, Invalidenstrasse 43, D-10115 Berlin, Germany. Sven O. Egenhoff Colorado State University, Department of Geosciences, 322 Natural Resources Building, Fort Collins, Colorado 80523-1482, USA
ABSTRACT Detailed mapping and facies analysis of a thick succession of diamictites of the Upper Ordovician Cancañiri Formation in southern Bolivia has revealed a glacioterrestrial origin for these sediments. The Cancañiri diamictites were deposited during three advances of a temperate, grounded ice sheet. They contain subglacial, englacial, and proglacial outwash sediments that increase in abundance from southeast to northwest. Clast fabrics and deformation features indicate SSE to NNW motion of the ice masses. Components of the diamictites usually display abrasion features such as facets and glacial striae. Provenance studies indicate that the pebbles comprise ~35% of siliciclastic sediments, mainly from the underlying shallow marine Ordovician rocks, 27% of slightly metamorphosed sediments that in part can be attributed to the Precambrian–Cambrian Puncoviscana Formation of northwestern Argentina, and a crystalline basement suite of metamorphic rocks (18%) and magmatic (mainly plutonic) rocks (20%). Due to the absence of typical lithologies, the Brazilian Shield, the Paraguay belt, and the southern Arequipa-Antofalla block could be excluded as possible source areas. The crystalline and metasedimentary clasts display strong affinities with the Pampean basement in central Argentina. All data consistently suggest that the Cancañiri tillites of southern Bolivia were deposited by a regional, lowlatitude ice sheet that was independent of the main inland ice mass of Gondwana and centered SSE of the study area, in a Neoproterozoic to Cambrian orogenic belt in the area of the present Argentinean Chaco. Keywords: Cancañiri diamictites, southern Bolivia, glacioterrestrial environment, Pampean provenance, South American ice sheet, Late Ordovician
*E-mail:
[email protected]. Schönian, F., and Egenhoff, S.O., 2007, A Late Ordovician ice sheet in South America: Evidence from the Cancañiri tillites, southern Bolivia, 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. 525–548, doi: 10.1130/2007.2423(27). For permission to copy, contact
[email protected]. ©2007 Geological Society of America. All rights reserved.
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INTRODUCTION The Hirnantian stage of the latest Ordovician saw the coincidence of the second largest mass extinction of the “big five” in the Phanerozoic (Sepkoski, 1996; Sheehan, 2001) and a major glaciation of the Gondwanan continent (Harland, 1972; Berry and Boucot, 1973; Hambrey, 1985; Ghienne, 2003) within a period of pCO2 levels sixteen times higher than at present (Crowley and Baum, 1991; Marshall et al. 1997; Kump et al., 1999; Berner and Kothavala, 2001). It is generally accepted that the main glaciation was restricted to the Hirnantian and had a short duration of only 0.5 to 1 Ma. It was preceded by a global warming event and accompanied by a glacio-eustatic sea level drop of ~50–100 m and positive isotopic excursions of δ18O, δ13C, and Δ13C, respectively (Brenchley et al., 1994, 2003; Marshall et al., 1997; Kump et al., 1999; Sheehan, 2001; Fortey and Cocks, 2005). Direct evidence of glaciation has been found in widespread glacial deposits in northern Africa that extend eastward to the Arabian peninsula and westward to Sierra Leone and Ghana (Biju-Duval et al., 1981; Deynoux and Trompette, 1981; Sutcliffe et al., 2000; Ghienne 2003; Le Heron et al., 2007; review in Sheehan, 2001). Farther south, glacioterrestrial deposits of the Late Ordovician belong to the Pakhuis Formation of the Table Mountain Group in South Africa (Rust, 1981; Hambrey, 1985; Sutcliffe et al., 2000), the Iapó tillites of the Paraná basin in southern Brazil (Rocha-Campos, 1981; Benedetto et al., 1992; Assine et al., 1994), and the Don Braulio Formation in the Precordillera of western Argentina (Buggisch and Astini, 1993; Astini, 1999). Reconstructions commonly display a single Late Ordovician ice sheet that covers the entire continent of Gondwana, with a south pole located in central–western Africa (e.g., Sutcliffe et al., 2000; Sheehan, 2001; “large scenario” of Ghienne, 2003). However, paleomagnetic data suggest a Late Ordovician south pole in northwestern Africa (Scotese et al., 1999; Li and Powell, 2001; Evans, 2003; McElhinny et al., 2003), which would be consistent with the main north African ice cap (Biju-Duval et al., 1981; Deynoux, 1985; Le Heron et al., 2005, 2007). The South American and South African tillites would then have been situated at considerably low paleolatitudes (30°–55°S; Evans, 2003), and extensive intracratonal areas of Gondwana at intermediate to low latitudes would have been covered with ice. Based on the apparent younger ages of glacial deposits in South America, Caputo and Crowell (1985) suggested that glacial centers migrated across Gondwana through the Late Ordovician and the Early Silurian (cf. Hambrey, 1985). Therefore, glacigenic deposits at an intermediate position between the Early Silurian diamictites reported from Brazil (Grahn and Caputo, 1992) and the Late Ordovician tillites in western Argentina, which have been attributed to local mountain glaciers (Buggisch and Astini, 1993; Astini, 1999; Peralta and Carter, 1999), are crucial for improving our knowledge of the early Paleozoic glacial period.
THE CANCAÑIRI (ZAPLA) FORMATION IN THE CENTRAL ANDES Diamictites close to the Ordovician-Silurian boundary in the Central Andes were first described from the Bolivian Altiplano within a mining report by Koeberling (1919) and named “Grauwacke de Cancañiri.” Later more systematic studies followed from southern Perú, from northwestern and central Bolivia, and from the so-called “Horizonte Glacial de Zapla” in northwestern Argentina (Schlagintweit, 1943; Davila and Ponce de Leon, 1971; Antelo, 1973, 1978; Anaya et al., 1987; Suárez Soruco, 1995; Boso, 1996; Díaz-Martinez et al., 2001; DíazMartinez and Grahn, 2007; Fig. 1A). Together with its lateral equivalent in Argentina, the Cancañiri Formation is exposed on an 1200-km long north-south transect covering an area that shows a maximum extension of more than 600 km in an eastwest direction from the sub-Andean ranges to the central Altiplano (Rodrigo et al., 1977; Suárez Soruco, 1977; Crowell et al., 1981). The thickness of the diamictite-bearing successions varies between less than 100 m in the eastern and southeastern realm (Fig. 1 A) and more than 1000 m in the Altiplano region (e.g., Tica-Tica locality in Fig. 1A). The origin of these formations has controversially been discussed in regional literature. Interpretations range from true tillites (Schlagintweit, 1943; Sempere, 1995; Martínez, 1998) over glaciomarine deposits with resedimentation features (Turner and Mendez, 1975; Crowell et al., 1981; Díaz-Martinez and Grahn, 2007) to debrites that have been resedimented during tropical climates (Anaya et al., 1987; Boso, 1996). That the diamictites of the Altiplano and adjacent Eastern Cordillera regions were deposited in a marine environment is evident from their great thickness, their abundant resedimentation features, and the presence of fossils in marine interbeds between diamictite units (Díaz-Martinez and Grahn, 2007). Striated clasts were attributed to local mountain glaciers and are thought to have been resedimented in a tectonically active basin (Rodrigo et al., 1977; Suárez Soruco, 1977; Sempere, 1995; Egenhoff, 2000). However, in contrast to the situation for central Bolivia and northwestern Argentina, few data have yet been published on the Cancañiri Formation of southern Bolivia (Fig. 1A; e.g., Schlagintweit, 1943; Suárez Soruco, 1977; Crowell et al., 1981). STUDY AREA The study area is situated 20 km north of Tarija and 5 km east of San Lorenzo between the villages of Sella Mendéz and Negro Muerto (lon 64°40′W, lat 21°23′S; Fig. 1A and B). While the Cancañiri Formation south of the study area is difficult to access and exposed as a narrow band of variable but low thickness (from 20 m at the Argentine-Bolivian border to 100 m in the Tarija region; Suárez Soruco, 1977; Crowell et al., 1981; Fig. 1A), it is this area where the formation is easily accessible and has been exposed several times due to intense west-verging Andean back-thrusting (Kley et al., 1996). Such outcrop
A Late Ordovician ice sheet in South America
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B
(L. Ordovician / Lower Arenig)
C northern cross-section
southern cross-section
Figure 1. (A) Generalized geologic map of the Central Andes, modified from Reutter et al. (1994), showing Neoproterozoic to Devonian sedimentary successions and the distribution and localities of the Cancañiri (Zapla) Formation. (B) Geologic map of a part of the study area with the location of the main sections and selected localities within the Cancañiri Formation. (C) Cross-sections displaying the west-verging backthrusting and the multiple exposure of the diamictites. Arr.—Arroyo; COL.—Colombia; PAR.—Paraguay; PT—Proterozoic; Qda.—Quebrada; Sra.—Sierra; UR.—Uruguay.
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conditions provide the opportunity to study lateral facies variations within the Cancañiri Formation from east to west and from south to north (Fig. 1B and C). Within the study area, the thickness of the succession of diamictites increases from 100 m in the south to >180 m in its northern parts (Figs. 1B and 2A). The diamictites overlie with a strong erosional unconformity a thick Lower Ordovician shallow marine succession of the Sella (>850 m) and Rumi Orkho (120–150 m) Formations (Schönian, 2003; Fig. 1B and C). This sequence contains abundant tempestites and indicators of a littoral, sub- to intertidal environment toward the top. It was deposited within a short time span in the lower Arenig on a shallow marine,
storm-dominated platform with deltaic influence (Egenhoff, 2000, 2003; Egenhoff et al., 2004). These sediments were uplifted and subaerially exposed during the late Early to early Middle Ordovician (Bahlburg and Hervé, 1997; Egenhoff, 2000, 2003). The diamictites are covered by one or two 2- to 15-m thick beds of shallow marine ironstones or ferriferous sandstones (Fig. 2A). The overlying 350- to 400-m thick succession of siltstones and sandstones of the Kirusillas Formation were deposited at shallow marine, low-energetic conditions, with an assumed coastline toward the west. The uppermost diamictites were reworked in a marine brackish to lacustrine environment after subaerial exposure. They contain chamosite ooids, indicating the presence of
Figure 2. (A) Image of Pucara Mountain from the southern Chaupi Cancha Valley to the northeast, with the Cancañiri Formation resting on the Rumi Orkho Quartzites and overlain by the ferriferous Kirusillas Sandstones (Quechua: “Pucara” = castle). Note the quartzitic sandstone bodies at the base and the continuous level of conglomeratic sandstones in the middle of the image. (B) Base of the Cancañiri Formation in the Chaupi Cancha Valley (Chaupi Cancha section, Fig. 4D), with irregular bodies of quartzitic sandstones. (C) Example of a sandstone lens with sharp contacts. (D) Detail of the same quartzitic sandstone showing its parallel lamination. (E) Comparison of thin sections (crossed nicols) of (bottom) a well-sorted quartzitic sandstone from the lower part of the formation and (top) a poorly sorted sandstone from the conglomeratic sandstone layer upsection (cf. Fig. 2A). FM—Formation.
A Late Ordovician ice sheet in South America a hiatus prior to the transgression of the ferriferous sandstones. Strongly altered feldspars within these diamictites and the occurrence of iron-rich deposits suggest a significant climatic change (Schönian, 2003). One east-west–striking fault within the valley of the Sella River and three north-south–striking faults can be identified in the study area (Fig. 1B). Among those, the very shallow El Potrero and Chaupi Cancha thrust faults account for a lateral displacement of ~0.8 and 1.4 km, respectively (Fig. 1B and C; Schönian, 2000). This implies that the outcrops of the diamictites were originally situated at considerably greater distances from each other and covered an area with an extension of around 4.5 instead of 2.3 km from east to west (Fig. 1B and C). In the Rumi Orkho Valley, only the upper half of the formation is exposed, where it is a relatively homogeneous succession of matrix-supported diamictites with only minor intercalations of quartzites or sandstones. Exposures on the western flank of the Chaupi Cancha Valley account only for the lower part of the formation (Fig. 2B). The complete succession can best be studied on the poorly vegetated hillsides of Pucara Mountain (Figs. 2A). FACIES ANALYSIS Standard lithofacies codes were introduced by Eyles et al. (1983) for the field description of (glacial) diamicton and diamictite successions and associated siliciclastic sediments. Following this classification, the Cancañiri Formation in southern Bolivia would be completely composed of massive, matrix-supported
A
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diamictites (“Dmm”) with some minor intercalations of sorted siliciclastics and clast-supported diamictites (Suárez Soruco, 1977; Crowell et al., 1981). For a more detailed vertical description of the succession, a lithofacies scheme based on the grain size classification of Moncrieff (1989) was developed that includes nine classes of diamictite lithofacies and twelve classes for better-sorted siliciclastics (Fig. 3; Schönian, 2000, 2003; Schönian et al., 2000). In addition to these lithofacies types, codes for grain size specifications (f~, m~, and g~ for fine, medium, and coarse) and for sedimentary structures, similar to those of Eyles et al. (1983), were used in the present study. Basal Contact The basal contact with the underlying Ordovician rocks is poorly exposed and could be studied at only four localities within the study area. In the Jarcas Valley in the south, the Cancañiri Formation overlies a 0.5- to 1-m thick transitional layer of heavily brecciated sandstones of the Rumi Orkho Formation. At the base of the Pucara section in the north, the Lower Ordovician siltstones, sandstones, and quartzites are followed up-section by 3 m of a monomict, internally deformed, mud-rich breccia. This chaotic deposit contains predominantly subsurface material, such as deformed silt- and sandstones and angular fragments of quartzites, but no exotic (i.e., far-traveled) lithologies (section A1 in Fig. 4A). In contrast, no deformation features could be observed at a basal contact in the Chaupi Cancha Valley (section E; Figs. 1B and 4A). Here the basal diamictites, which appear
B
Figure 3. (A) Petrographic triangle for poorly sorted sedimentary rocks, modified from Moncrieff (1989). The boundary between diamictites and conglomerates was defined as 70% gravel content according to Füchtbauer (1988), and additional classes for clast-rich diamictites with a gravel content of 40–70% were introduced. (B) The nine classes of grain size distributrions obtained for diamictites and the twelve for sorted siliciclastics as well as the thirteen codes used to specify sedimentary characteristics can be easily assigned in the field.
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Figure 4. (A) Sections of the base of the Cancañiri Formation taken east (right) and west (left) of the Chaupi Cancha thrust fault. Note the softsediment deformation features at the base of section A1, which are absent from the fossil-bearing, bioturbated succession west of the Chaupi Cancha thrust. (B–D) Vertical sections measured in the study area. The Jarcas section (B) is situated in the southeast, the Pucara section (C) in the north (both east of the Chaupi Cancha thrust; Fig. 1B), and the Chaupi Cancha section (D) in the northwest (Fig. 2B). Classes of grain size distributions and codes used to specify sedimentary characteristics as given in Figure 3B.
A Late Ordovician ice sheet in South America to have been slumped, conformably overlie an 8- to 9-m thick marine sequence of siltstones and fine sandstones that contain bivalves, brachiopod fragments, and trace fossils. Similarly, the base of the Chaupi Cancha section farther north does not display evidence of subsurface deformation and is marked by a 1.5- to 2-m thick intercalation of bioturbated siltstones and quartzitic sandstones that disconformably rest on top of the Rumi Orkho Formation (Fig. 4D). Vertical Sections Three vertical sections were measured in the study area, one in the Jarcas Valley of the Negro Muerto area, one on the northern side of Pucara Mountain, and one along the northwestern flank of the Chaupi Cancha Valley (Figs. 1B and 4). In addition to the basal marine intercalation west of the Chaupi Cancha Fault (member A), three members were identified within the massive succession of diamictites (members B, C, and D). They are separated by levels of laterally continuous bodies of sandstones, quartzites, and conglomerates that occur at approximately one-third and twothirds of the formation up-succession (55–77 m and 105–127 m above the base) and are usually marked by a sharp contact at their top (Figs. 2A and 4). The Jarcas section starts with the heterogeneous member B (Fig. 4A), which in the lower 50 m is composed of massive diamictites with an intermediate matrix type and rare clasts (I[D]– ID). This lower part contains abundant lenticular bodies of very well-sorted, parallel-laminated to cross-stratified quartzitic sandstones that reach thicknesses of up to 10 m and are laterally continuous over more than 30 m. The diamictites of the upper part of member B are more heterogeneous in matrix composition and clast content (Fig. 4B). Toward the top of the member, only small bodies of sorted siliciclastics and irregularly stratified diamictites occur. The uppermost meters are characterized by a stratified succession of coarse, gravelly sandstones, sandy mudstones with rare clasts, and a continuous layer of conglomeratic diamictite with a gravellag on its surface (m76; Fig. 4B). Member C starts with a sandy diamictite with internal shearing features. It shows a gradual change from intermediate to sandy diamictites from base to top (Fig. 4B). The diamictites are again more heterogeneous up-succession, but intercalations of well-sorted clastics are rare. The upper 20 m are represented by a massive body of intermediate to sandy diamictites, which at its base and top grades into coarse sandstones with rare clasts and is bound by sharp contacts. Member D is relatively homogeneous and, except for a finer, internally sheared lower part, is composed of coarse diamictites with a sandy matrix. The Pucara section most clearly reveals the threefold succession of the Cancañiri diamictites (Fig. 4C). Above the basal deformed breccia, fine-grained, heterogeneous diamictites with intercalated muddy sandstones are exposed in the lower 20 m of the succession (section A1 in Fig. 4A and C). They are overlain by 10 m of relatively homogeneous intermediate to sandy diamictites. Well-sorted quartzitic sandstones are absent. However, as in
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the Jarcas section, the succession becomes more heterogeneous and contains small sandstone bodies and irregularly stratified diamictites in the upper part of the member. Its top is marked by a thick, laterally constant body of conglomeratic sandstone or fine conglomerate (Fig. 4C). Member C is very similar to member B, starting with finer, sheared diamictites and passing toward a heterogeneous succession with abundant lenticular bodies of sorted siliciclastics and irregularly stratified diamictites. It is again topped by a laterally continuous, albeit thinner, layer of fine conglomerate (Fig. 4C). The upper member D is thicker than in the Jarcas section, but as well composed of homogeneous, coarse diamictites with an intermediate to sandy matrix type. The Chaupi Cancha section (Fig. 4D), located west of the Chaupi Cancha Thrust at a paleodistance of ~2–2.5 km from the Pucara locality, comprises mainly member B. The lowermost diamictites are very fine-grained and clast-poor and overlain by relatively heterogeneous intermediate to sandy diamctites with numerous intercalations of lenticular sandstone bodies (Figs. 2B and 4D). As in the Pucara section, the middle part is composed of heterogeneous intermediate to sandy diamictites, but contains abundant intercalations of sorted siliciclastics. Toward the top (m42–m61), the succession passes from nonstratified, massive deposits to well-stratified and better-sorted, clast-supported diamictites ([S]D) and conglomerates that often show gradual contacts with each other. Intercalated bodies of conglomerates and conglomeratic sandstones as well as gravellags are abundant (Fig. 4D). The top of member B is here characterized by a laminated, in part cross-stratified, graded, gravelly sandstone bed, which is again separated by a sharp contact from the overlying intermediate diamictites of member C. Associated Facies Elements, Vertical Trends, and Lateral Variations Due to the small-scale lateral and vertical variability of the Cancañiri Formation, often gradational contacts, and strong weathering in part, it is difficult to define single facies elements within the massive diamictites. Nevertheless, typical facies associations and facies trends can be recognized. Fine-grained, often rather homogeneous diamictites that in part display internal shear features do occur at the base of each of the three members of the diamictite succession. In the Jarcas and Chaupi Cancha Valleys, they are usually associated with lenticular bodies of laminated to cross-stratified quartzitic sandstones that display sharp contacts with the confining diamictites. These sandstones are well sorted (sorting after Inman: 0.6–1.2; Füchtbauer, 1988), display a skewness of +0.05 to +0.35, are physically mature, and are composed mainly of quartz grains with only rare feldspars, rock fragments, and diamictite pellets (Fig. 2B–E). Coarse, clast-rich diamictites in the upper part of member B and C are associated with siliciclastics that often show irregular geometries, may display sharp or gradual contacts with the diamictites, and cover a wider spectrum of grain sizes and variable degrees of sorting (0.8–2.5) and skewness (+0.15 to +0.57; cf. Fig. 2E). They are
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Figure 5. (A) Field image of the northern Pucara Mountain, the conglomeratic sandstone layer at the top of member B, and members C and D of the Pucara section. (B) Detailed sections measured along this layer. Note its heterogeneity and the irregular lower and very sharp upper contact. (C) Deformed fine-grained, clast-poor diamictite below a medium, sandy conglomerate in the Chaupi Cancha Valley (DP1 in Fig. 1B). Deformation occurred from southeast to northwest. Classes of grain size distributions and codes used to specify sedimentary characteristics as given in Figure 3B. FM—Formation.
Angular clasts of quartzites and sandstones from the Rumi Orkho Formation that document subsurface erosion occur only in the lower ~5–10 m of the succession (Figs. 4A and 6A). At one locality in the Chaupi Cancha Valley, a large block of quartzitic sandstone with a prominent set of striations in a SSE to NNW direction was found within the lowermost fine-grained diamictites (M[D]; Fig. 6B). Up-section clasts are usually better rounded and display evidence of abrasion, such as facets or striae (Fig. 6C–E). To infer the transport path of the clasts, more than three hundred pebbles were collected from all stratigraphic levels of the Cancañiri Formation. They all have been cleaned carefully and analyzed for abrasion features. This analysis revealed that >80% of the clasts display at least one faceted surface and that around 65% have facets on more than one side. Similarly, >60% of them are striated, with around 30% showing striae in two or more directions (Fig. 7A). There is no major variation in the number of faceted and striated clasts from the base to the top of the succession. Both harder lithologies, such as granites and quartzites, and softer ones, such as mudstones, display facets and striae (Fig. 6C–E). Some of the pebbles resemble glacigenic “stoss-lee” clasts (Fig. 6C; see Boulton, 1978). This, together with the high number of faceted and striated clasts and the occurrence of typical “nailhead striations” on some of them (Fig. 6E), indicates a glacial origin of these abrasion features. To statistically assess the degree of abrasion, the roundness of completely preserved (n = 263) clasts was determined and their sphericity was calculated (using the roundness scale and “intercept sphericity” of Krumbein, 1941). Figure 7B demonstrates that both roundness and sphericity are variable and again without any specific trend throughout the succession. Clast roundness spans values from 0.3 to 0.9 (subangular to well rounded), with a maximum at around 0.55 (subrounded;
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Figure 6. (A) Brecciation, deformation, and erosion at the base of the Jarcas section in the southeastern study area within a brittle shear zone (glaciotectonite). (Fig. 1B). Note the angular blocks of quartzite in the basal diamictites. (B) Parautochthonous block of fossiliferous, quartzitic sandstone in the basal diamictites with two sets of striae (close to section E in Fig. 1B). The predominant striation is oriented SSE to NNW. (C) “Stoss-lee” clast of fine sandstone with multiple sets of striae from the matrix of the lower Cancañiri Formation (DP2 in Fig. 1B). The truncated lee side faced toward WNW, while the main striation was roughly oriented SSE to NNW. (D) Strongly faceted coarse-grained granite clast (aplitic granite, Fig. 11J). (E) Mudstone with two sets of deeply grooved striae (left) and quartzitic clasts displaying the glacial “nailhead striation” (right).
Fig. 7B). Clasts are mainly spheroidal and oblate-spheroidal (planar). Prolate-spheroidal (tallud) types of clasts are less abundant, and triaxial (planar-tallud) clasts are very rare. In consequence, the sphericity is very high and falls between the values of 0.55 and 0.95 (Fig. 7B). Where possible, the long axes of clasts were measured in situ within the diamictites. The resulting fabrics show a relatively chaotic pattern for all values, but a consistent SSE to NNW orientation for clasts from member B (Fig. 7C). The latter suggests that the deviations observed for the complete succession are due either to chaotic fabrics in its upper part or to systematically distinct orientations within each of members B, C, and D. Facies Interpretation A glacioterrestrial origin has been suggested by Schönian et al. (1999) for the diamictites of the study area, thus indicating
that most of them represent true tillites in the strict sense (Boulton and Deynoux, 1981; Deynoux, 1985; Dreimanis and Schlüchter, 1985; Dreimanis, 1989). The main lines of evidence are: (1) the erosional basal contact with shearing, glacial abrasion, and soft sediment deformation of the underlying Lower Ordovician rocks, represented by brittle glaciotectonites (Fig. 6A) and deformation tillites (Fig. 4A) in the eastern study area; (2) the evidence of reworking within an intertidal to brackish environment atop the undulating surface of the uppermost diamictites prior to the transgression of the basal-Silurian ferriferous sandstones (Schönian, 2003); (3) the association of polymict, mostly matrix-supported diamictites with channelized sandstones and conglomerates and, except for the base of the formation in the western study area (Fig. 4A), the lack of any marine intercalation within the diamictites; and (4) the shape, the degree of roundness, and the glacial abrasion features of both locally eroded as well as far-traveled clasts, which can be related to processes in the basal
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Faceted clasts
Faceted and Striated Clasts
Faceted clasts
Faceted clasts
Faceted clasts
Figure 7. (A) Abundance of roundness of the pebbles and percentage of clasts that exhibit facets and striae for the complete succession and each of the three members (grayscale of the lower bars indicates the number of clasts with facets on two, three, or four sides or striations in two or three directions). (B) Sphericity/roundness diagrams of pebbles collected from the different members (grayscale represents 100%, 85%, and 50% of the clasts for each group). (C) Rose diagram of the fabrics of the long axis of elongate clasts measured in the field (values rotated by general dip). Fm.—Formation; Mbr.—Member.
zone of traction of a glacier (e.g., Kjaer, 1999). The abundance of these features excludes a possible resedimentation of material from adjacent local mountain glaciers that should contain supraglacially derived, less abraded debris from high-level transport. In turn, the high degree of roundness and sphericity as well as the abrasion features of the clasts imply that most, if not all, of them have undergone at least one phase of low-level transport. This, together with the apparent nivelation of a pre-existing relief on a supraregional scale, is typical of an inland ice sheet (Boulton, 1978; Eyles et al., 1983; Miller, 1996; Benn and Evans, 1998). The diamictites of the study area as well as the overlying Kirusillas Formation do not display the large-scale marine
resedimentation features that have been reported for the marine Cancañiri Formation in central Bolivia (Díaz-Martínez and Grahn, 2007). Glaciogenic resedimentation seems to be restricted to the supra- and proglacial deposits separating the individual units, and the Kirusillas ironstones, sandstones, siltstones, and shales have most likely been deposited in a relatively quiet shelf environment. In the absence of striated pavements, the extent of past ice sheets may be reconstructed with the help of subglacial deformation features (e.g., Le Heron et al., 2007). For the Cancañiri Formation of the study area, a transition from ice contact facies to a floating ice sheet across the Chaupi Cancha thrust fault is evident for basal member A, but fully glacioterrestrial conditions have
A Late Ordovician ice sheet in South America been established during the deposition of members B, C, and D (discussed later). The detailed petrographic description and the observed internal structures of the diamictites were used to attribute lithofacies types to one of the end-members of till-forming processes (deformation, lodgement, or melting; “till triplet” after Boulton and Deynoux, 1981; “till tetrahedon” after Dreimanis, 1989; “spaghetti diagrams” after Benn and Evans, 1998). Deformation tills contain deformed subsurface substrate and/or internal deformation features. Lodgement tills are clast poor and have a fine-grained matrix, may display joints, and are usually stacked as “till sheets,” whereas melt-out tills are coarser, display preserved internal structures, and show well-developed clast fabrics. Because till-forming processes form a continuum, Benn and Evans (1998) used the terms “till evolution” for the processes related to till formation and “hybrid tillites” for the stacked sequence of polygenetic lodgement and melt-out tills that results from a repeated decoupling of the glaciers from their base during phases of ice stagnation. The geometry of the sandstone bodies within the Cancañiri Formation; their sedimentary structures, such as lamination and cross-stratification; the positive skewness of their grain size distribution; and their absence of fossils indicate a (glacio-) fluvial origin of these deposits (Füchtbauer, 1988). The lenses of wellsorted sandstones observed in the lower part of members B and C were most likely formed subglacially, whereas the less wellsorted conglomerates and gravelly sandstones in the upper part of these members represent supra- to proglacial outwash (Jurgaitis and Juozapavicˇius, 1989; Lawson, 1989). By combining the inferred till-forming and secondary processes, the position of the material within the glacier during deposition could be determined and the associated lithofacies could be assigned to the respective depositional environment (Brodzikowski and van Loon, 1987; Miller, 1996; Benn and Evans, 1998). Following the glacial tectonites (Jarcas section), assimilated deformation tillites, and partially assimilated deformation tillites (Pucara section) at the base of member B, its lower part is composed of polygenetic, lodgement, and melt-out tillites that were deposited at the base of the glacier (to m50 in the Jarcas section, to m32 in the Jarcas, and to m30 in the Chaupi Cancha section; Fig. 4). Lodgement appears to be predominant in the southeast, whereas the importance of melting seems to increase toward the northwest. The well-sorted, lenticular sandstones indicate a highly energetic flow in the upper flow regime within outwash channels. These deposits of the subglacial environment are followed up-section by a transition zone with probably englacial sediments and then by supraglacial melt-out tillites (from m62, m43, and m42, respectively). The clast-supported diamictites, conglomerates, and pebbly sandstones at the top of member B represent diamictite material reworked either by debris flows or by localized streams at different flow regimes. These supraglacial sediments pass to laterally constant, finer, and better-sorted proglacial outwash deposits in the northern and northwestern parts of the study area (Figs. 4 and 5A–B; Brodzikowski and van Loon, 1987; Jurgaitis and Juozapavicˇius, 1989; Lawson, 1989).
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The sharp upper contact of these sediments can be attributed to subglacial erosion by the second ice advance, which is most likely also responsible for the soft-sediment deformation within the upper part of member B observed in the northern Chaupi Cancha Valley (Fig. 5C). A similar sequence of subglacial, englacial, supraglacial, and proglacial sediments was deposited in member C, albeit with less abundant and thinner outwash deposits, which suggests a more proximal environment. Member D is almost entirely composed of melt-out tillites, probably deposited in an englacial to supraglacial environment. Supraglacial or proglacial outwash deposits are absent, maybe because of erosion during subsequent subaerial exposure (Fig. 4). The massive and thick, matrix-supported lodgement tillites in members B and C that are associated with channelized outwash, the abundant resedimentation deposits, and the large amount of glacially abraded debris indicate a temperate (warmbased or wet-based) thermal regime for the ice sheet (Eyles et al., 1983; Miller, 1996; Benn and Evans, 1998). PROVENANCE OF THE CANCAÑIRI TILLITES A qualitative and semiquantitative analysis of the pebbles was performed in order to characterize the source area of the Cancañiri Formation. All identifiable clasts have been classified lithologically (n = 284; Fig. 8A). The spectra obtained were corrected for small- to medium-sized quartz or quartzite clasts (n = 93; 44%), which could not be determined macroscopically. Without these lithologies, the pebbles of the Cancañiri Formation include ~35% siliclastic, slightly diagenetically altered rocks; 27% highly diagenetic or slightly metamorphosed siliciclastic deposits; 10% metasediments and parametamorphic rocks; 8% orthometamorphites; >16% plutonites; and around 2–5% volcanic rocks. Their relative abundance varies throughout the succession, with the highest number of sedimentary clasts in member B and the quantity of crystalline clasts increasing toward the top (Fig. 8A). The roundness-sphericity diagram suggests that abrasion is largely independent of lithology (Fig. 8B). However, a greater variability in roundness and sphericity for sedimentary than for metamorphic or magmatic clasts can be observed. Thin sections of sixty-six of the clasts were analyzed for a detailed petrographic description. The analyzed clasts are weakly diagenetized sediments (twenty-three thin sections), strongly diagenetized or metamorphically overprinted sediments (nine thin sections), metasedimentary clasts (eight thin sections), orthometamorphites (twelve thin sections), and magmatites (fourteen thin sections). Sedimentary Clasts Clasts with a Low Degree of Diagenesis The sedimentary clasts, mainly collected from member B, usually contain fossils. Two clasts with Cruziana trace fossils, which are abundant in the Rumi Orkho Formation underlying the diamictites, have been found (Fig. 9A). Many clasts contain
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paleotaxodont and/or ctenodont bivalves (Fig. 9B–C) and therefore cannot be older than the Early Ordovician. One block of fine-grained sandstone shows flaser bedding and bioturbation features and bears a coquina layer of bivalves (Fig. 9B). The parautochthonous, quartzitic sandstone block of Figure 6B contains ostracodes and tentaculites, as well as trilobite, brachiopod, and recrystallized bivalve fragments, filaments, and phosphate pellets (Fig. 9D). Many clasts indicate a highly variable intertidal environment, such as a fossil-rich sandstone; a tempestitic, fossilrich, fine-grained sandstone with pelitic intraclasts and glauconite; a flat-pebble conglomerate; laminated and bioturbated mudstones; and black, phosphatic sandstones (Fig. 9D–H). These
clasts display a mixed metamorphic and plutonic provenance, and volcanic or volcaniclastic components are absent. Well-Diagenetized Clasts This group includes all clasts that show a high degree of diagenesis and often a slight metamorphic (or tectonic) overprint. They are transitional to the metasediments (Fig. 10). Such clasts occur throughout the formation and are not predominant in any member. They are highly variable siliciclastic deposits of different degrees of physical and chemical maturity. These sediments fall into three categories: fine-grained, immature silt- or sandstones of magmatic and metamorphic provenance and a possible volcanic contribution; poorly sorted quartzitic arkoses, graywackes, and diamictites with variable amounts of volcanic components; and relatively mature quartzitic arkoses and sandstones (Fig. 10A–F). The latter display a mixed plutonic and metamorphic provenance. Metamorphic Clasts Metasediments The metamorphic clasts include clasts of a sedimentary origin that display clear evidences of metamorphic overprint. Metamorphism is usually low to medium grade (epizone, zeolith facies to lower greenschist facies) and pressure induced. Petrographically, these rocks resemble the higher diagenetic siliciclastics and include immature, mica-rich metaquartzites, low-grade metamorphic arkoses and graywackes, and, to a lesser degree, highly mature metaquartzites (Fig. 10H–L). One strongly abraded and weathered clast of a low-grade, volcaniclastic graywacke was found within the study area (Fig. 10G). Again, the poorly sorted rocks seem to contain volcanic or volcaniclastic components (Fig. 10H–J), whereas the chemically and physically mature clasts have a cratonic provenance (plutonic or metamorphic quartzes). The metapelite of Figure 10L has a preserved original lamination, but is strongly chloritized, contains abundant calcite veins, and can be classified as hornfels gneiss.
ñ
Figure 8. (A) Quantitative classification of the clasts of the Cancañiri Formation. Magm.—magmatites; Metamorph.—metamorphites; Ortho—orthometamorphites; Para—parametamorphites; Plut.—plutonites; sd Sed.—slightly diagenetized sediments; Volc.—volcanites; wd—well diagenetized. (B) Sphericity/roundness diagrams for different lithologies (grayscale represents 100%, 85%, and 50% of the clasts).
Orthometamorphites The degree of metamorphism among these clasts ranges from low-grade plutonic lithologies (metagranites, -granodiorites, and -syenites; Fig. 11A–F) over medium-grade orthogneisses of mesozone (pressure-induced) metamorphism, greenschist- and amphibolite-facies rocks to high-grade granulitic lithologies with at least partial anatexis (Fig. 11C). While a barrow-type metamorphism is predominant, some clasts seem to have suffered relatively high temperatures (contact and regional metamorphism). One very weathered, andesitic volcanic clast with a serial porphyric texture and a slight metamorphic overprint has been found (Fig. 11F). However, because volcanic rock fragments are present within the diamictite matrix and in sedimentary and metasedimentary clasts, the record seems to be biased by weathering. Therefore, the Cancañiri Formation of the study area (Fig. 8A) is estimated to contain 2–5% volcanic clasts.
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Figure 9. Sedimentary clasts with a low degree of diagenesis. (A) Medium sandstone with the trace fossil Cruziana, cf. rugosa (m70, Jarcas section). (B) Block of bioturbated fine sandstone with flaser bedding and a coquina layer (m47, Pucara section). (C) Bivalves from sandstone clasts (various localities within member B). (D) Scan (left) of thin section (plane light) and microphotograph (right) of the striated block of fine, laminated, quartzitic sandstone of Figure 6B with phosphatized bioclasts of tentaculites, brachiopodes, trilobites, phosphate pellets, and quartzitic recrystallized bivalves. (E) Fossil-rich siltstone (left) with intraclasts of poorly sorted mudstones (tempestite). Bioclasts are phosphatized brachiopods, bivalves, and trilobites. Intraclasts contain quartz and glauconite (right, arrow) within a muddy or phosphatic matrix. (F) Soft pebble conglomerate of elongated clasts of laminated mudstones weakly cemented with carbonate and phosphate (m34, Pucara section). (G) Heavily bioturbated, laminated siltstone mudstone (m9, Chaupi Cancha section). (H) Thin section microphotograph of a medium sandstone with phosphate (black), carbonate cement, and a phosphatized bioclast (left). All thin section microphotographs under crossed polarized light. Scale bar, if not indicated differently, is 100 µm.
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Figure 10. Well-diagenetized to tectonically overprinted sedimentary clasts and low- to medium-grade metasediments. (A) Prolate-spheroid clasts of green, quartzitic sandstone with striations on the faceted side, parallel to its long axis. (B) Scan of (left) a thin section of laminated siltstone with (right) microfractures. (C) Unsorted graywacke or diamictite with secondary sericite in the matrix. Zonar plagioclase and sanidin indicate a volcaniclastic contribution. (D) Tectonically overprinted, immature feldspar sandstone (lithic arkose) with weathered (volcanic?) rock fragments and idiomorph plagioclase crystals. (E) Coarse, quartzitic arkose, metamorphically overprinted, with weathered rock fragments. (F) Quartzite with still distinguishable grain boundaries and cement, but dynamic recrystallization along weakened zones. (G) Striated, low-grade metamorphic, mica-rich, volcaniclastic graywacke (left) with idiomorph sanidine, zonar plagioclase, and weathered anorthoclase (right) within a matrix of sericite with preferred orientation. (H) Low-grade metamorphic, matrix-supported subarkose or graywacke with neoblastic quartz in pressure shadows and a recrystallized matrix of oriented sericite and muscovite. (I) Low-grade metagraywacke with albitized plagioclase and beginning chloritization of mica. (J) Low-grade paragneiss, feldspar- and mica-rich, with beginning chloritization. (K) Low-grade, very mature metaquartzite. (L) Green metapelite with chlorite (lower right) and calcite (left) veins (hornfels gneiss). All thin section microphotographs under crossed polarized light. Scale bar, if not indicated differently, is 100 µm.
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Figure 11. Orthometamorphic and magmatic clasts. (A) Feldspar-rich, two-mica gneiss (orthogneiss, mesozone). (B) Green, chloritic orthogneiss, with rare pyrite (upper greenschist facies). (C) Garnet-bearing, felsic granulite gneiss (upper amphibolite to granulite facies). (D) Metamorphically overprinted granodiorite with beginning chloritization of mica (lower greenschist facies). (E) Orthogneiss with granophyric structures and quartz neoblasts (metasubvolcanite). (F) Scan of (left) low-grade metamorphic, andesitic volcanite with serial porphyric texture that contains (right) mainly plagioclase and newly formed muscovite, biotite, and rare chlorite. (G) Heavily weathered, coarse, equigranular, aplitic granite, Jarcas Valley. (H) Granite within a diamictite matrix at a fresh river cut, northern Chaupi Cancha Valley. (I) Tectonically deformed, mica-poor, tourmaline-bearing granite. (J) Deformed, mica-poor, zircon-bearing aplitic granite (Fig. 6D). (K) Tectonically deformed (low-grade metamorphic), two-mica granodiorite. (L) Deformed, late magmatic granodiorite, with rare myrmekite. (M) Pegmatitically overprinted leucogranite, only weakly deformed. All thin section microphotographs under crossed polarized light. Scale bar, if not indicated differently, is 100 µm.
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Magmatic (Plutonic) Clasts Plutonic lithologies often weather rapidly, which sometimes makes it difficult to determine the amount of feldspars and mica. Identified magmatic clasts (including tectonized lithologies) comprise aplites, granites, granodiorites, diorites, and monzodiorites (Fig. 11G–M). The largest clast found in the study area (1.7 m; Fig. 11G) is a coarse-grained, mica-poor aplitic granite. One clast exhibits large prismatic zircons, and another is rich in turmaline (Fig. 11H and I). All of the plutonic clasts analyzed show a very low-grade metamorphic overprint or are at least deformed. Some of the orthometamorphic and plutonic lithologies are pegmatitic or have a pegmatitic or late magmatic overprint (Fig. 11E and L).
crystalline clasts, comprises a continental basement of acidic to intermediate chemistry with highly variable metamorphic rocks and an orogenic suite of plutonites and subvolcanites that were affected by tectonic overprint or low-grade metamorphism after emplacement (Fig. 11). The provenance of the sand fraction of the diamictites and associated siliciclastics is in good agreement with the results from clast analysis, but displays a slight chemical fractionation toward granitic feldspars by ice transport and a strong physical fractionation toward stable quartzes and heavy minerals by glaciofluvial reworking of the diamictites. DISCUSSION AND CONCLUSIONS Possible Source Areas
Provenance of the Diamictites and Sandstones Rock fragments within the diamictite resemble the lithologies found among the clasts of the Cancañiri Formation. Feldspar grains, which make up 10–30% of the sand fraction, are mainly plutonic potassium feldspars (microcline, perthites, microperthites). Sand-sized plagioclase fragments, abundant in crystalline clasts and rock fragments, are very rare in the diamictites and associated sandstones, but do occur again in their fine silt fraction. This fractionation is probably related to plagioclase instability during glacial transport, which is documented for Pleistocene ice sheets (Füchtbauer, 1988). Volcanic or metamorphic feldspars (albite, sanidine) are rare. Cherts and chertoid grains are probably derived from metapelitic or volcanic rocks. The bulk of the sand grains are mono- and polyquartzes. About 30% of the quartzes are homogeneous, often contain inclusions of heavy minerals, and indicate a plutonic origin. Some 20% are tectonically overprinted and display a slight undulosity. The rest of the quartz grains show a highly variable grade of metamorphism and are characterized by deformation bands, formation of subgrains, or dynamic recrystallization. The heavy minerals that occur within the diamictite matrix are, in descending abundance, zircons, tourmalines of variable colors, epidote, glaucophane, and garnets. The abundance of sedimentary clasts in the lower succession suggests that a large part of them was eroded from underlying rocks that were later covered by glacial deposits (Fig. 8A). Their variable roundness and sphericity (Fig. 8B) might be explained by the occurrence of subangular fragments derived directly from the subsurface and of softer sedimentary pebbles that, when transported over considerable distances, abrade rapidly. The group of fossil-rich clasts of a shallow marine environment (Fig. 9) was most likely eroded within this region from the underlying Rumi Orkho Formation (Egenhoff, 2000, 2003; Egenhoff et al., 2004). Their cratonic provenance is consistent with a source in the Brazilian Shield to the east as it was inferred for the Lower Ordovician of southern Bolivia (Egenhoff and Lucassen, 2003). Farther away, presumably older, poorly sorted siltstones, arkoses, and graywackes and low-grade volcanic or volcaniclastic lithologies were exposed, and the mature sandstones and metaquartzites probably represent a third source of siliciclastic rocks (Fig. 10). The most distant area, as characterized by the
Because the area of the present Altiplano formed a tectonically active marine basin throughout the Late Ordovician and earliest Silurian (Rodrigo et al., 1977; Crowell et al., 1981; Sempere, 1995; Suárez Soruco, 1995; Díaz-Martinez et al., 2001; Díaz-Martinez and Grahn, 2007), no glaciers can have reached southern Bolivia from the northwest. Thus, four basement areas are discussed as possible sources for the Cancañiri tillites of the study area and are outlined in their paleogeographic context (the roman numerals refer to the areas shown in Fig. 12A). I. The southwestern Amazonian craton This is a highly heterogeneous basement complex with an accretion history from the Paleoproterozoic Transamazonian orogenic cycle to the Neoproterozoic pan-African orogenesis (BritoNeves et al., 1999; Sial et al., 1999; Almeida et al., 2000; Loewy et al., 2004). The discussed area embraces the Paleoproterozoic to Mesoproterozoic Juruena province, the Mesoproterozoic Rondonia province, the Neoproterozoic Sunsas province (Guaporé craton) that extends to eastern Bolivia, and the Tucavaca belt of the Brasilides (Fig. 12A). The complex Juruena province is in part highly metamorphic (amphibolite to granulite facies) and is composed of granitic-migmatitic complexes, granulites, and tonalitic gneisses as well as anorogenic Rapakivi-type granites, syenites, gabbros, mangerites, and charnockites. Paleoproterozoic metavolcanic and volcaniclastic series including banded iron formations are exposed (Tassinari and Macambira, 1999). The Rondonia province is similar, but additionally contains archaic complexes of gabbroid-volcaniclastic units and greenstone belts as well as rhyodacitic and trachytic lavas, basaltic dikes, and very mature siliciclastic molasse deposits of the Sunsas cycle (Saes and Leite, 1993; Tassinari and Macambira, 1999). The Sunsas province displays a similar basement and includes metamorphosed Mesoproterozoic sedimentary series as psammitic to pelitic schists, meta-arkoses, quartzites, basaltic metavolcanites, calcitic paragneisses, and metaironstones that are associated with mafic to ultramafic volcanic rocks (Litherland and Power, 1989). The latter also accompany the Neoproterozoic sandstones, arkoses, quartzites, pelitic rocks, and iron-rich sediments of the Sunsas cycle and the sediments of the pan-African Tucavaca belt, which include diamictites, graywackes,
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A
B
Figure 12. (A) The possible source areas for the Cancañiri Formation of the study area discussed in the text. (B) The Pampean Massif below the southern Chaco of northwestern and central Argentina, the windows into its basement, and Neoproterozoic to Upper Ordovician orogenic belts. Note that after the Middle Ordovician accretion of the Precordillera terrane and the Famatina orogeny, and after the Late Ordovician Ocloyic orogeny in northwestern Argentina, the proto-Andean margin of Gondwana was largely accreted as it is today (see text). Modified from Goudarzi (1977).
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cherts, banded ironstones, dolomites, and limestones (Litherland and Bloomfield, 1981; Brito-Neves et al., 1999).
formation is folded and to the west unconformably overlain by the Cambrian Mesón Group.
II. The Paraguay belt This region east of the study area forms part of the Brasilides and can be divided into a Transamazonian basement block (Rio Apa craton; Fig. 12A and B) and the pan-African fold belt. The former is composed of the medium- to high-grade metamorphic Rio Apa Complex represented by orthogneisses, garnet-mica schists, paragneisses, quartzites, and amphibolites and by the felsic tuffs, volcanites, and late to post-tectonic granites of the Aluminador Suite and the Amonquija Group (Brito-Neves et al., 1999). The Paraguay fold belt, sensu strictu, includes an eastern internal basement complex where Cambrian granites, para- and orthogneisses, and Neoproterozoic tonalites, granodiorites, rhyolites, and metavolcanites are exposed and a western external zone composed of sediments of the Sunsas cycle (diamictites, conglomerates, graywackes, arkoses, arenites with mangan and iron concretions, phosphates, limestones, dolomites, sandstones, red shales, and siltites) that are progressively metamorphosed toward the internal zone (Alvarenga and Trompette, 1993; Brito-Neves et al., 1999; Sial et al., 1999).
IV. The proto-Andean margin This region southwest of the study area is composed of Precambrian medium- to high-grade metamorphic basement of the southern Arequipa-Antofalla block and the mobile Puncoviscana, Famatina, and Ocloyic belts of northwestern Argentina (Breitkreuz et al., 1989; Reutter et al., 1994; Bahlburg and Hervé, 1997; Saavedra et al., 1998; Keppie and Bahlburg, 1999; Lucassen et al., 2000; Loewy et al., 2004; Fig. 12). The former is composed of metasediments (banded paragneisses, mica schists and gneisses, amphibolites), Ordovician orthometamorphites (orthogneisses, migmatites, anatectic granites, metabasalts) of Mesoproterozoic protoliths, and Carboniferous metaquartzites, mica schists, amphibolites, and metabasalts with a Proterozoic protolith age (Baeza and Pichowiak, 1988; cf. Reutter et al., 1994; Loewy et al., 2004). The orogenic belts SSW of the study area include highly heterogeneous crystalline basement complexes of Precambrian to Upper Ordovician age that are in the north covered by the Vendian–Cambrian Puncoviscana Formation and the Cambrian Mesón Group and are toward the south progressively more highly metamorphosed due to the later Variscan to Andean exhumation (Saavedra et al., 1998). In the north, granitoid intrusions of a problematic Ordovician age are reported from the “Faja Eruptiva de la Puna Oriental” and the Pampean sedimentary series were intruded by the Middle Cambrian Cañani and Santa Rosa de Tastil Granites (granodiorites, granites, tonalites), but are covered unconformably by Upper Cambrian to Lower Ordovician sedimentary rocks (Bahlburg and Hervé, 1997; Lucassen et al., 2000). The most striking characteristics of the Puna region are the thick Lower to Middle Ordovician acidic to mafic volcaniclastic series derived from the Puna arc to the west (debrites, tuffs, rhyolitic ashs, mafic lavas, pillow basalts, hydroclastites, turbidites, volcaniclastic conglomerates, sandstones, siltstones), which were intruded locally by Lower Ordovician plutons (Breitkreuz et al., 1989; cf. Reutter et al., 1994; Lucassen et al., 2000). The preservation of these series and evidences of an Early Silurian transgression in the Puna region (the “Llandovery beds” at the Salar de Rincón; Bahlburg and Hervé, 1997) indicate that the uplift of the Puna region by the Late Ordovician Ocloyic orogeny was not very pronounced and that more deeply seated, sediment-covered basement rocks were probably not exhumed by that time (cf. Turner and Mendez, 1975; Sempere, 1995; Bahlburg and Hervé, 1997). Mafic to ultramafic rocks, Rapakivi-type granites, and many of the volcaniclastic, sedimentary, and metasedimentary lithologies of the Amazonian craton have not been found within the Cancañiri Formation of southern Bolivia. Many of the sedimentary, metamorphic, and magmatic rocks exposed in the Paraguay belt to the east are absent from the diamictites and, due to the scarcity of volcanic and volcaniclastic clasts and the abundance of metamorphic and plutonic lithologies, provenance from the uplifted, volcanic Arequipa-Puna region is highly unlikely. The
III. The Pampean Massif This massif or terrane to the southeast is covered by Phanerozoic sediments. The only windows into its basement are the Rio Apa craton (described earlier) and the eastern Sierras Pampeanas (Sierras de Cordoba and Sierra Norte Ambargasta; Fig. 12A and B). The oldest unit of the latter consists of Precambrian sedimentary rocks of highly variable metamorphism including mature red quartzarenites, quartzitic sandstones, and sandstones with calcite and hematite cements, as well as metamorphic semipelitic quartz-feldspar schists and quartzites. The polymetamorphic crystalline basement is composed of greenschist- to granulite-facies rocks including phyllites, schists, metapelites, metapsammites, para- and orthogneisses, various migmatites, and rare granulitic rocks (Baldo et al., 1996; Lira et al., 1997; Rapela et al., 1998; Schwartz and Gromet, 2004). The magmatic suite of the eastern Sierras Pampeanas is related to the pan-African and Pampean orogeny (600–490 Ma) and is represented by partly porphyritic granodiorites and monzogranites, biotite orthogneisses and tonalites, dioritic-tonalitic gneisses, pegmatitic and migmatitic granites, leucogranites, and cordieritites, as well as trondhjemites, leucotonalites, aplogranites, and postorogenic subvolcanites with porphyric or granophyric textures (Lira et al., 1997; Rapela et al., 1998). In the north, the Pampean basement is most likely covered by an eastern equivalent of the Vendian to Lower Cambrian Puncoviscana Formation, which is a succession composed of low-grade metamorphic, immature diamictites, graywackes, arkoses, sand- and siltstones with associated conglomerates, red or green pelitic rocks, volcanic lavas and dikes, and rarely limestones in its upper parts (Aceñolaza et al., 1988; Bahlburg and Hervé, 1997; Keppie and Bahlburg, 1999; Schwartz and Gromet, 2004; Fig. 12B). The
A Late Ordovician ice sheet in South America metasedimentary and crystalline basement clasts of the diamictites match best with the Pampean basement reported from the Sierra de Cordoba and Sierra Norte Ambargasta. The assemblage of immature, poorly sorted, and low-grade metamorphic siltstones, sandstones, graywackes, and diamictites with partially volcanic provenance and the rare volcanic and volcaniclastic lithologies resembles the Neoproterozoic to Lower Cambrian Puncoviscana Formation of northwestern Argentina (Fig. 12B). Transport Direction Favorable outcrop conditions in an easily accessible area and a newly developed scheme of lithofacies codes made it possible to trace lateral facies variations within the massive glacigenic Cancañiri diamictites of southern Bolivia with hitherto unknown detail. Sedimentary features together with provenance data consistently suggest an ice transport from the southeast or SSE toward the northwest or NNW for this region. A summary of the principal pieces of evidence includes the following: 1. the increasing thickness of the diamictite succession toward the glaciomarine basin northwest of the study area (Figs. 1 and 4); 2. pronounced brittle shearing and subglacial soft-sediment deformation at the base of member B, in the southeast and east of the study area, which are absent west of the Chaupi Cancha thrust fault (Figs. 4A and 6A; cf. Fig. 1B), indicating a transition from an erosive grounded to a partially floating ice sheet; 3. the increasing abundance of englacial, supraglacial, and proglacial outwash deposits and resedimentation features in members B and C, suggesting transition from a proximal to a more distal environment from the southeast to the northwest (Figs. 4 and 5A–B); 4. motion indicators such as intraformational deformation features (Fig. 5C), striations on parautochthonous blocks from the underlying Ordovician (Fig. 6B), and faceted and striated “stoss-lee” clasts (Fig. 6C); 5. the flow-parallel alignment of the long axes of clasts within the massive tillites in a SSE to NNW direction (~160° SE or 340° NW; Fig. 7C); and 6. probable provenance from a metamorphic or plutonic crystalline basement area in the Pampean Massif to the southeast to SSE (Pampean orogen?) that was partially covered by a low-grade metasedimentary turbiditic sequence with volcanic influence to the SSE (Puncoviscana Formation, Fig. 12B). In conclusion, the Cancañiri Formation near Sella was deposited during three advances of a temperate ice sheet from the SSE. There is evidence of subglacial deformation and erosion at the base of member B that diminishes to the northwest and therefore indicates a transitional position toward a glaciomarine environment beyond the grounding line of the ice sheet. The second and third ice advances left the former deposits largely intact and stacked an unusually thick sequence of terrestrial tillites.
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Age and Correlations The lower Paleozoic diamictites in Bolivia and Argentina were originally considered of Silurian age (Branisa et al., 1972; Antelo, 1973, 1978; Turner and Mendez, 1975; Crowell et al., 1981). No biostratigraphically indicative fossils have yet been reported from southern Bolivia, neither from within the Cancañiri Formation nor from the Silurian shallow marine Kirusillas Formation. The biostratigraphic data from central Bolivia are controversial, for while Branisa et al. (1972) recognized the Wenlock acritarch Leifusa within the Cancañiri Formation, their fauna from the overlying Sacta Limestone has a Llandovery to early Wenlock age (Díaz-Martinez, 2007). Similarly, the “Cancañiri Fauna” of Antelo (1973) was actually derived from alternating silt- and sandstones above the diamictites in the Cochabamba area. Late Ordovician trilobites and the brachipod Hirnantia cf. sagittifera from within the formation (Anaya et al., 1987; Benedetto et al., 1992; Toro, 1994), as well as Early Ashgill brachiopods and graptolites from below the diamictites (Suárez Soruco and Benedetto, 1996; Toro and Salguero, 1996), confine its deposition to a timespan from the Middle Ashgill to the Late Llandovery. In northwestern Argentina, Ashgillian trilobites were found within the Zapla Formation in the Sierra de Santa Barbara (southeastern part of Fig. 1A; Monaldi and Boso, 1987). Additionally, graptolites from between the first and second ferriferous sandstone beds that overly the diamictites in the “Puesto Viejo” mine revealed an age close to the Ordovician-Silurian boundary (persculptus or atavus biozone; Boso and Monaldi, 1987; Monteros et al., 1993; Fig. 1A). The best stratigraphic control for early Paleozoic glacial diamictites in the Andes is that of the lower member of the Don Braulio Formation of the Argentine Precordillera, which is confined to the Upper Ashgill (Buggisch and Astini, 1993; Astini, 1999; Fig. 13A). The marine deposits immediately above the diamictites provided a typical Hirnantia fauna, and the overlying siltstones of the upper Don Braulio Formation yielded the latest Ordovician graptolite Glyptograptus persculptus (Benedetto, 1986; Sanchez et al., 1991; Marshall et al., 1997; Peralta and Carter, 1999). Above the G. persculptus occurrences, these siltstones contain, much as in northwestern Argentina, oolitic ironstones at the base of the Silurian (atavus biozone). Lithologically, the Cancañiri and Kirusillas Formations of southern Bolivia resemble the Zapla and Lipeon Formations of northwestern Argentina rather than the marine successions across the Ordovician-Silurian boundary of central Bolivia, where they have been formally defined (Branisa et al., 1972; Antelo, 1978; Benedetto et al., 1992; Monteros et al., 1993). Stratabound oolitic ironstones or ferriferous sandstones in the lowermost Llandovery, which are probably related to transgressive ravinement surfaces that truncated Fe-saturated estuaries (Boso and Monaldi, 1987; Astini et al., 2003), have no equivalent in central Bolivia and can not be correlated with the much younger Sacta Limestone Member following the Cancañiri diamictites in the Cochabamba area (Díaz-Martínez, 2007). The occurrence of 100 m of marine siltand sandstones separating the uppermost diamictites from the
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A
B PROTO-ALPINE TERRANES
- South pole - Paleolatitudes
Figure 13. (A) Reconstruction of the “Pampean ice sheet” in South America with Hirnantian tillites, the hypothetic glacial center, and inferred transport directions (not palinspastically restored; base map as in Fig. 12B). The area south of the Ocloyic belt was probably elevated due to the repeated mountain-building throughout the early Paleozoic. (B) Gondwana in the latest Ordovician with the locations of the assumed ice sheets and areas with a questionable or unknown record (paleogeography from Scotese et al., 1999; directions of ice transport from Ghienne, 2003). The pole position is a compromise between the 450 Ma paleomagnetic South Pole offshore northwest Africa of Li and Powell (2001) and the 455 Ma pole of McElhinny et al. (2003) some 5°–10° farther inland.
ironstones in the “Las Pircas” mine (Fig. 1A; Boso and Monaldi, 1987) is consistent with the inferred hiatus atop the uppermost diamictites in the study area. Thus, the diamictites in southern Bolivia and northwestern Argentina are directly correlatable with one another and with the Hirnantian lower member of the Don Braulio Formation and were most probably deposited prior to the G. persculptus biozone (Monteros et al., 1993). A Pampean Ice Sheet in the Late Ordovician The diamictites of the Don Braulio Formation were interpreted as terrestrial lodgement and melt-out tillites of temperate (warm- or wet-based), grounded mountain glaciers that show intercalations of glacioaquatic deposits (Buggisch and Astini, 1993; Astini, 1999; Peralta and Carter, 1999). Three separate ice advances are reported, and a grooved pavement indicates a transport direction from ESE to WNW (Astini, 1999; Fig. 13A). Similarly, the Zapla Formation in northwestern Argentina seems to represent a threefold succession. A glacially striated surface that suggests sediment transport from the southeast to the northwest was observed at a locality at the Argentine-Bolivian border, only ~80 km south of the study area (Martínez, 1998; Santa Victoria Range; Fig. 13A; “Qda. Puestos” locality in Fig. 1A). Some authors report an eastward continuation of the
Zapla Formation in the subsurface of the present Chaco region, where it unconformably covers Cambrian and Lower Ordovician rocks and reaches thicknesses of up to 100 m (Benedetto et al., 1992; Franca et al., 1995). Based on the interpretations of the southern Bolivian Cancañiri Formation presented here; the inferred transport direction, which shows a strong deviation from an expected movement away from the pole (Fig. 13); and the correlatable three ice advances and postglacial iron-rich deposits, we suggest that all the diamictites in the southern Central Andes were deposited by one regional, low-latitude, temperate ice sheet. This independent South American ice sheet was probably centered in the Pampean orogen in the region of the present southern Chaco of central–northern Argentina (Rapela et al., 1998), covered the complete Pampean Massif, and extended toward the orogenic belts of the proto-Andean margin of Gondwana (Fig. 13A). In the eastern Paraná basin, the Upper Ordovician Iapó diamictites have been interpreted as terrestrial tillites (Rocha-Campos, 1981; Benedetto et al., 1992; Assine et al., 1994). However, a marine succession across the Ordovician-Silurian boundary without evidence of glacigenic deposits in the western Paraná basin indicates that the Pampean glaciation did not extend toward central Gondwana (Benedetto et al., 1992; Assine et al., 1994; Fig. 13A). It is likely that the Iapó tillites represent
A Late Ordovician ice sheet in South America deposits of another ice sheet and can be related to the South African Pakhuis glaciation (Fig. 13B). These arguments favor a “multiple ice-sheet scenario” for the Late Ordovician ice age, which would significantly reduce the area of Gondwana covered by ice (the “small scenario” of Ghienne, 2003; Fig. 13B) and may also explain the less pronounced glacioeustatic regression (<100 m) compared to the Pleistocene (Brenchley et al., 1994, 2003; Sheehan, 2001). Such a scenario would alleviate the effects of positive albedo feedback and inhibited silicate weathering in the paleoclimatic and paleo-oceanographic models applied to explain the apparent paradox of a sudden glaciation within a greenhouse period (Crowley and Baum, 1991; Brenchley et al., 1994, 2003; Kump et al., 1999; Gibbs et al., 2000; Sutcliffe et al., 2000; Herrmann et al., 2004). Recently discussions suggested that the early Paleozoic glacial period might have already started in the late Middle Ordovician and lasted well into the Silurian (Azmy et al., 1998; Pope and Read, 1998; Herrmann et al., 2004; Saltzman and Young, 2005; cf. Hambrey, 1985). The Hirnantian age of the Pampean ice sheet and its position at latitudes as low as 35°–45° (or 45°– 55° according to McElhinny et al., 2003; Fig. 13B) question the hypothesis of migrating glacial centers and focus the discussion on Early Silurian glaciations toward diamictites of a probable Llandovery–Wenlock age from the Amazonas basin in Brasil (Grahn and Caputo, 1992). Evidence of grounded glaciers was also found at exposures of the Cancañiri Formation in central Bolivia (e.g., the “roches moutonnées” at the Pojo locality; Branisa et al., 1972). From north of Cochabamba, fossiliferous, shallow marine limestones were reported that are over- and underlain by diamictites, which might indicate two different glaciation events (Anaya et al., 1987; Perez Guarachi, 1989; Díaz-Martinez, 2007). However, at present it is not clear whether Late Ordovician and/or Early Silurian glaciers reached Bolivia from the Amazonian craton to the northeast (Fig. 13B). A more detailed basin analysis and better biostratigraphic control of the well-exposed Cancañiri Formation in Bolivia would help to reconstruct the areal extent of the ice sheet(s), improve the stratigraphic timing of the glaciation(s), and provide field data for paleoclimatologic modeling, and would thus significantly enhance our understanding of the dramatic environmental changes that affected the Earth during this period. ACKNOWLEDGMENTS FS dedicates this paper to his daughter, Maria Gabrielle Alvarado Pena. We thank Bernd-D. Erdtmann (Technical University, Berlin), Joachim Marcinek (Humboldt–University of Berlin), and Bernd Weber (Free University, Berlin) for their supervision, field assistance, sedimentological and paleontological advice, and fruitful discussions. Peter Brauner, Cornelia von Engelhardt, Hanne Glowa, Bernd Kleeberg, and Silke Stöwer are gratefully acknowledged for their logistic support
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and assistance in laboratory work. Martin Harnisch and Friedrich Lucassen contributed to the determination of crystalline lithologies and to discussions on the basement of the South American platform. Enrique Díaz-Martinez, Jonas Kley, Oliver Lehnert, and Ramiro Suárez Soruco provided rare literature on the regional geology. FS is very grateful for the support of the citizens of Sella Mendez during his stay, especially to Mario, the teachers of the local school, and Doña Petrona, and he wants to thank the people of Tarija, who helped in many ways, especially José Paz Garcón and family, and Lilo and Carlos Methfessel, Ursula Tarraga-Wiedemann, and Juan Tarraga. Many thanks to Jörg Maletz for his revision of the manuscript and to Pat J. Brenchley, Petr Štorch, and Andrey Zhuravlev for their constructive reviews, which helped to improve the article considerably. Fieldwork was supported by the Collaborative Research Center 267 (“Deformation Processes in the Andes”) of the German Science Foundation (DFG, SFB 267, project C3). REFERENCES CITED Aceñolaza, F.G., Miller, H., and Toselli, A.J., 1988, The Puncoviscana Formation (Late Precambrian Early Cambrian)—Sedimentology, tectonometamorphic history and age of the oldest rocks of northwestern Argentina, in Bahlburg, H., et al., eds., The Southern Central Andes: Lecture Notes in Earth Sciences, v. 17: Berlin, Springer, p. 25–37. Almeida, F.F.M., Brito-Neves, B.B., and Carneiro, C.D.R., 2000, The origin and evolution of the South American Platform: Earth-Science Reviews, v. 50, p. 77–111. Alvarenga, C.J.S., and Trompette, R., 1993, Evolucao tectônica brasiliana da faixa Paraguai: A estruturacão da Regão de Cuiabá: Revista Brasileira de Geosciencias, v. 23, no. 1, p. 18–30. Anaya, F., Pacheco, J., and Perez, H., 1987, Estudio estratigráfico-paleontológico de la Formación Cancañiri en la Cordillera del Tunari (Departamento de Cochabamba), in IV Congreso Latinoamericano de Paleontología, Santa Cruz, Bolivia, Memorias, p. 679–693. Antelo, B., 1973, La Fauna de la Formación Cancañiri (Silúrico) en los Andes centrales bolivianos: Revista del Museo de La Plata, v. 7, no. 45, p. 267–277. Antelo, B., 1978, Las formaciones de edad Silúrica en el noroeste argentino (provincias de Jujuy y Salta): Revista de la Asociacion Geológica de Argentina, v. 33, no. 1, p. 1–16. Assine, M.L., Soares, P.C., and Milani, E.J., 1994, Seqüências tectono-sedimentares mesopaleozóicas da bacia do Paraná, Sul do Brasil: Revista Brasileira de Geociências, v. 24, no. 2, p. 77–89. Astini, R.A., 1999, The Late Ordovician glaciation in the Proto-Andean margin of Gondwana revisited: Geodynamic implications: Acta Universitatis Carolinae–Geologica, v. 43, nos. 1–2, p. 171–173. Astini, R.A., Marengo, L., and Rubinstein, C.V., 2003, The Ordovician stratigraphy of the Sierras Subandinas (Subandean Ranges) in Northwest Argentina and its bearing on an integrated foreland basin model for the Ordovician of the Central Andean Region, in Albanesi, G.L., et al., eds., Ordovician from the Andes: INSUGEO, Serie Correlación Geológica, v. 17: Tucuman, Argentina, p. 381–386. Azmy, K., Veizer, J., Basset, M.G., and Copper, P., 1998, Oxygen and carbon isotopic composition of Silurian brachiopods: Implications for coeval seawater and glaciations: Geological Society of America Bulletin, v. 110, no. 11, p. 1499– 1512, doi: 10.1130/0016-7606(1998)110<1499:OACICO>2.3.CO;2. Baeza, L., and Pichowiak, S., 1988, Ancient crystalline basement provinces in the north Chilean Central Andes: Relics of continental crust development since the Mid Proterozoic, in Bahlburg, H., Breitkreuz, Ch., and Giese, P., eds., The Southern Central Andes: Lecture Notes in Earth Sciences, v. 17: Berlin, Springer, p. 3–24. Bahlburg, H., and Hervé, F., 1997, Geodynamic evolution and tectonostratigraphic terranes of northwestern Argentina and northern Chile: Geological
546
Schönian and Egenhoff
Society of America Bulletin, v. 109, no. 7, p. 869–884, doi: 10.1130/00167606(1997)109<0869:GEATTO>2.3.CO;2. Baldo, E.G., Demange, M., and Martino, R.D., 1996, Evolution of the Sierras de Córdoba, Argentina: Tectonophysics, v. 267, p. 121–142, doi: 10.1016/ S0040-1951(96)00092-3. Benedetto, J.L., 1986, The first typical Hirnantia fauna from South America (San Juan Province, Argentine Precordillera), in Racheboef, R.R., and Eming, C.C., eds., Les brachiopodes fossiles et actuels: Biostratigr. Paléozoïque, v. 4, p. 439–447. Benedetto, J.L., Sanchez, T.M., and Brussa, E.D., 1992, Las cuencas Silúricas de América Latina, in Gutiérez Marco, F.G., et al., eds., El Paleozoico Inferior de Ibero-América: Cáceres, Spain, Universidad de Extremadura, p. 119–148. Benn, D.I., and Evans, D.J.A., 1998, Sediment facies, in Glaciers and glaciation: London, Arnold, p. 377–420. Berner, R.A., and Kothavala, Z., 2001, GEOCARB III: A revised model of atmospheric CO2 over Phanerozoic time: American Journal of Science, v. 301, p. 182–204. Berry, W.B.N., and Boucot, A.J., 1973, Glacio-eustatic control of Late Ordovician–Early Silurian platform sedimentation and faunal changes: Geological Society of America Bulletin, v. 84, p. 275–284, doi: 10.1130/00167606(1973)84<275:GCOLOS>2.0.CO;2. Biju-Duval, B., Deynoux, M., and Rognon, P., 1981, Late Ordovician tillites of the Central Sahara, in Hambrey, M.J., and Harland, W.B., eds., Earth’s pre-Pleistocene glacial record New York, Cambridge University Press, p. 99–107. Boso, M.A., 1996, Los ambientes sedimentarios de la formación Zapla en el noroeste argentino: XII Congreso Geológico de Bolivia, Tarija, Memorias, p. 951–960. Boso, M.A., and Monaldi, C.R., 1987, Oolitic stratabound iron ores in the Silurian of Argentina and Bolivia, in Fontboté, L., et al., eds., Stratabound ore deposits in the Andes: Berlin, Springer, p. 175–186. Boulton, G.S., 1978, Boulder shape and grain-size distribution of debris as indicators of transport paths through a glacier and till genesis: Sedimentology, v. 25, p. 773–799, doi: 10.1111/j.1365-3091.1978.tb00329.x. Boulton, G.S., and Deynoux, M., 1981, Sedimentation in glacial environments and the identification of tills and tillites in ancient sedimentary sequences: Precambrian Research, v. 15, p. 397–420, doi: 10.1016/03019268(81)90059-0. Branisa, L., Chamot, G.A., Berry, W.B.N., and Boucot, A.J., 1972, Silurian of Bolivia, in Berry, W.B.N., and Boucot, A.J., eds., Correlation of South American Silurian rocks: Boulder, Colorado, Geological Society of America Special Paper 133, p. 21–31. Breitkreuz, C., Bahlburg, H., Delakowitz, B., and Pichowiak, S., 1989, Paleozoic volcanic events in the Central Andes: Journal of South American Earth Sciences, v. 2, no. 2, p. 171–189, doi: 10.1016/0895-9811(89)90045-X. Brenchley, P.J., Marshall, J.D., Carden, G.A.F., Robertson, D.B.R., Long, D.G.F., Meidla, T., Hints, L., and Anderson, T.F., 1994, Bathymetric and isotopic evidence for a short-lived Late Ordovician glaciation in a greenhouse period: Geology, v. 22, p. 295–298, doi: 10.1130/00917613(1994)022<0295:BAIEFA>2.3.CO;2. Brenchley, P.J., Carden, G.A., Hints, L., Kaljo, D., Marshall, J.D., Martma, J.D., Meidla, T., and Nõlvak, J., 2003, High-resolution stable isotope stratigraphy of Upper Ordovician sequence: Constraints on the timing of bioevents and environmental changes associated with mass extinction and glaciation: Geological Society of America Bulletin, v. 115, p. 89–104. Brito-Neves, B.B., Campos Neto, M.d.C., and Fuck, R.A., 1999, From Rodinia to Western Gondwana: An approach to the Brasiliano-Pan African Cycle and orogenic collage: Episodes, v. 22, no. 3, p. 155–166. Brodzikowski, K., and van Loon, A.J., 1987, A systematic classification of glacial and periglacial environments, facies and deposits: Earth-Science Reviews, v. 24, p. 297–381, doi: 10.1016/0012-8252(87)90061-4. Buggisch, W., and Astini, R., 1993, The Late Ordovician ice age: New evidence from the Argentine Precordillera, in Findlay et al., eds., Gondwana Eight: Assembly, evolution, and dispersal: Rotterdam, Balkema, p. 439–447. Caputo, M.V., and Crowell, J.C., 1985, Migration of glacial centers across Gondwana during the Paleozoic Era: Geological Society of America Bulletin, v. 96, p. 1020–1036, doi: 10.1130/0016-7606(1985)96<1020: MOGCAG>2.0.CO;2. Crowell, J.C., Suárez-Soruco, R., and Rocha-Campos, A.C., 1981, The Silurian Cancañiri (Zapla) Formation of Bolivia, Argentina and Peru, in Hambrey,
M.J., and Harland, W.B., eds., Earth´s pre-Pleistocene glacial record: Cambridge, England, Cambridge University Press, p. 902–907. Crowley, T.J., and Baum, S.K., 1991, Towards reconciliation of Late Ordovician (440 Ma) glaciation with very high CO2 levels: Journal of Geophysical Research, v. 96, p. 8793–8800. Davila, J., and Ponce de Leon, V., 1971, La sección del Río Inambiri en la faja subandina del Perú y la presencia de sedimentitas de la Formación Cancañiri (Zapla): Revista Técnica de Yacimientos Petroliferos Fiscales Bolivianos, v. 1, no. 1, p. 67–85. Deynoux, M., 1985, Terrestrial or waterlain glacial diamictites?: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 51, p. 97–141, doi: 10.1016/0031-0182(85)90082-3. Deynoux, M., and Trompette, R., 1981, Late Ordovician tillites of the Taoudeni Basin, West Afrika, in Hambrey, M.J., and Harland, W.B., eds., Earth’s pre-Pleistocene glacial record: Cambridge, England, Cambridge University Press, p. 89–98. Díaz-Martínez, E., 2007, The Sacta Limestone Member (early Wenlock): Coolwater, temperate carbonate deposition in the distal foreland of Gondwana’s active margin, Bolivia: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 245, no. 1-2, p. 46-61, doi: 10.1016/j.palaeo.2006.02.028. Díaz-Martínez, E., and Grahn, Y., 2007, Early Silurian glaciation along the western margin of Gondwana (Peru, Bolivia and northern Argentina): Palaeogeographic and geodynamic setting: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 245, no. 1-2, p. 62-81, doi: 10.1016/j.palaeo.2006.02.018. Díaz-Martinez, E., Acosta, H., Cardenas, J., Carlotto, V., and Rodriguez, R., 2001, Paleozoic diamictites in the Peruvian altiplano: Evidence and tectonic implications: Journal of South American Earth Sciences, v. 14, no. 6, p. 587–592, doi: 10.1016/S0895-9811(01)00054-2. Dreimanis, A., 1989, Tills: Their genetic terminology and classification, in Goldthwait, R.P., and Matsch, C.L., eds., Genetic classification of glacigenic deposits: Final report of the International Union of Quaternary Research (INQUA): Rotterdam, Balkema, p. 17–83. Dreimanis, A., and Schlüchter, C., 1985, Field criteria for the recognition of till or tillite: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 51, p. 7–14, doi: 10.1016/0031-0182(85)90079-3. Egenhoff, S.O., 2000, Sedimentologie und Beckenentwicklung im Ordovizium in Südbolivien: Berliner Geowissenschaftliche Abhandlungen, Reihe A, Band 207, 173 p. Egenhoff, S.O., 2003, Litoestratigrafía y sedimentología del Ordovícico Inferior del sur de Bolivia: Revista Técnica de Yacimientos Petroliferos Fiscales Bolivianos, v. 21, p. 147–156. Egenhoff, S.O., and Lucassen, F., 2003, Chemical and isotopic composition of Lower to Upper Ordovician sedimentary rocks (Central Andes / South Bolivia): Implications for their source: Journal of Geology, v. 111, p. 487– 497, doi: 10.1086/375280. Egenhoff, S.O., Maletz, J., and Erdtmann, B.-D., 2004, Lower Ordovician graptolite biozonation and lithofacies of southern Bolivia: Relevance for palaeogeographic interpretations: Geological Magazine, v. 141, p. 287– 299, doi: 10.1017/S0016756804009239. Evans, D.A.D., 2003, A fundamental Precambrian–Phanerozoic shift in Earth’s glacial style?: Tectonophysics, v. 375, p. 353–385, doi: 10.1016/S00401951(03)00345-7. Eyles, N., Eyles, C.H., and Miall, A.D., 1983, Lithofacies types and vertical profile models: An alternative approach to the description and environmental interpretation of glacial diamict and diamictite sequences: Sedimentology, v. 30, p. 393–410, doi: 10.1111/j.1365-3091.1983.tb00679.x. Fortey, A., and Cocks, L.R.M., 2005, Late Ordovician global warming: The Boda event: Geology, v. 33, no. 5, p. 405–408, doi: 10.1130/G21180.1. Franca, A.B., Milani, E.J., Schneider, R.L., and López, O.P., 1995, Phanerozoic correlation in southern South America, in Tankard, A.J., Suárez Soruco, R., and Welsink, H.J., eds., Petroleum basins of South America: Tulsa, Oklahoma, American Association of Petroleum Geologists Memoir 62, p. 129–161. Füchtbauer, H., 1988, Sedimente und Sedimentgesteine: Sedimentpetrologie, Teil 2: Stuttgart, Schweitzerbartsche Verlagsbuchhandlung, 1141 p. Ghienne, J.-F., 2003, Late Ordovician sedimentary environments, glacial cycles, and post-glacial transgression in the Taoudeni Basin, West Africa: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 189, p. 117–145, doi: 10.1016/S0031-0182(02)00635-1. Gibbs, M.T., Bice, K.L., Barron, E.J., and Kump, L.R., 2000, Glaciation in the Early Paleozoic “greenhouse”: The roles of paleogeography and
A Late Ordovician ice sheet in South America atmospheric CO2, in Huber, B.T., et al., eds., Warm climates in Earth history: Cambridge, England, Cambridge University Press, p. 386– 422. Goudarzi, G.H., 1977, Geological map of South America: Arlington, Virginia, U.S. Geological Survey, Miscellaneous Field Studies MF-868 A, scale 1:5,000,000. Grahn, Y., and Caputo, M.V., 1992, Early Silurian glaciations in Brazil: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 99, p. 9–15, doi: 10.1016/0031-0182(92)90003-N. Hambrey, M.J., 1985, The Late Ordovician–Early Silurian glacial period: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 51, p. 273–289, doi: 10.1016/0031-0182(85)90089-6. Harland, W.B., 1972, The Ordovician ice age: Geological Magazine, v. 109, p. 451–456. Herrmann, A.D., Haupt, B.J., Patzkowsky, M.E., Seidov, D., and Slingerland, R.L., 2004, Response of Late Ordovician paleoceanography to changes in sea level, continental drift, and atmospheric pCO2: Potential causes for longterm cooling and glaciation: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 210, p. 385–401, doi: 10.1016/j.palaeo.2004.02.034. Jurgaitis, A., and Juozapavicˇius, G., 1989, Genetic classification of glaciofluvial deposits and criteria for their recognition, in Goldthwait, R.P., and Matsch, C.L., eds., Genetic classification of glacigenic deposits: Final report of the International Union of Quaternary Research (INQUA): Rotterdam, Balkema, 1989, p. 227–241. Keppie, J.D., and Bahlburg, H., 1999, Puncoviscana Formation of northwestern and central Argentina: Passive margin or foreland basin deposit? in Ramos, V.A., and Keppie, J.D., eds., Laurentia-Gondwana connections before Pangea: Boulder, Colorado, Geological Society of America Special Paper 336, p. 139–144. Kjaer, K.H., 1999, Mode of subglacial transport deduced from till properties, Mýrdalsjökull, Iceland: Sedimentary Geology, v. 128, p. 271–292, doi: 10.1016/S0037-0738(99)00074-3. Kley, J., Gangui, A.H., and Krüger, D., 1996, Basement-involved blind thrusting in the eastern cordillera oriental, southern Bolivia: Evidence from cross-sectional balancing, gravimetric and magnetotelluric data: Tectonophysics, v. 259, p. 171–184, doi: 10.1016/0040-1951(95)00067-4. Koeberling, F.R., 1919, Informe sobre la geología de las propiedades mineras de la Compañia Estañifera de Llallagua: Informe inédito: Llallagua, Bolivia, Compania Estañifera de Llallagua. Krumbein, W.C., 1941, Roundness and shape of sedimentary particles: Journal of Sedimentary Petrology, v. 11, no. 2, p. 73–79. Kump, L.R., Arthur, M.A., Patzkowsky, M.E., Gibbs, M.T., Pinkus, D.S., and Sheehan, P.M., 1999, A weathering hypothesis for glaciation at high atmospheric pCO2 during the late Ordovician: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 152, p. 173–187, doi: 10.1016/S00310182(99)00046-2. Lawson, D.E., 1989, Glacigenic resedimentation: Classification concepts and application to mass-movement processes and deposits, in Goldthwait, R.P., and Matsch, C.L., eds., Genetic classification of glacigenic deposits: Final report of the International Union of Quaternary Research (INQUA): Rotterdam, Balkema, p. 147–169. Le Heron, D.P., Sutcliffe, O.E., Whittington, R.J., and Craig, J., 2005, The origin of glacially related soft-sediment deformation structures in Upper Ordovician glacigenic rocks: Implications for ice-sheet dynamics: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 218, p. 75–103, doi: 10.1016/j.palaeo.2004.12.007. Le Heron, D.P., Ghienne, J.-F., El Houicha, M., Khoukhi, Y., and Rubino, J.-L., 2007, Maximum extent of ice sheets in Morocco during the Late Ordovician glaciation: Palaeogeographic and geodynamic setting: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 245, nos. 1–2, p. 200–226, doi: 10.1016/j.palaeo.2006.02.031. Li, Z.X., and Powell, C.M., 2001, An outline of the palaeogeographic evolution of the Australasian region since the beginning of the Neoproterozoic: Earth-Science Reviews, v. 53, p. 237–277, doi: 10.1016/S00128252(00)00021-0. Lira, R., Millone, H.A., Kirschbaum, A.M., and Moreno, R.S., 1997, Calc-alkaline arc granitoid activity in the Sierra Norte-Ambargasta ranges, central Argentina: Journal of South American Earth Sciences, v. 10, no. 2, p. 157–177, doi: 10.1016/S0895-9811(97)00013-8. Litherland, M., and Bloomfield, K., 1981, The Proterozoic history of Eastern Bolivia: Precambrian Research, v. 15, p. 157–179, doi: 10.1016/03019268(81)90027-9.
547
Litherland, M., and Power, G., 1989, The geologic and geomorphologic evolution of Serranía Huanchaca, eastern Bolivia: The legendary Lost World: Journal of South American Earth Sciences, v. 2, no. 1, p. 1–17, doi: 10.1016/0895-9811(89)90023-0. Loewy, S.L., Connelly, J.N., and Dalziel, I.W.D., 2004, An orphaned basement block: The Arequipa-Antofalla basement of the Central Andean margin of South America: Geology, v. 116, nos. 1–2, p. 171–187. Lucassen, F., Becchio, R., Wilke, H.G., Franz, G., Thirlwall, M.F., Viramonte, J., and Wemmer, K., 2000, Proterozoic–paleozoic development of the basement of the Central Andes (18–26°s)—A mobile belt of the South American Craton: Journal of South American Earth Sciences, v. 13, p. 697–715, doi: 10.1016/S0895-9811(00)00057-2. Marshall, J.D., Brenchley, P.J., Mason, P., Wolff, G.A., Astini, R.A., Hints, L., and Meidla, T., 1997, Global carbon isotopic events associated with mass extinction and glaciation in the Late Ordovician: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 132, p. 195–210, doi: 10.1016/S00310182(97)00063-1. Martínez, M., 1998, Late Ordovician glacial deposits of northwest Argentina: New evidence from the Mecoyita Formation, Santa Victoria Range: Journal of African Earth Sciences, v. 27, no. 1A, p. 136–137. McElhinny, M.W., Powell, C.M., and Pisarevsky, S.A., 2003, Paleozoic terranes of eastern Australia and the drift history of Gondwana: Tectonophysics, v. 362, p. 41–65, doi: 10.1016/S0040-1951(02)00630-3. Miller, J.M.G., 1996, Glacial sediments, in Reading, H.G., ed., Sedimentary environments: Processes, facies and stratigraphy (3rd edition): Oxford, Blackwell, p. 454–484. Monaldi, C.R., and Boso, M.A., 1987, Dalmanitina (Dalmanitina) subandina nov. sp. (Trilobita) en la formación Zapla del norte argentino: IV Congreso Latinoamericano de Paleontología, Bolivia, Actas I, p. 149–157. Moncrieff, A.C.M., 1989, Classification of poorly sorted sedimentary rocks: Sedimentary Geology, v. 65, p. 191–194, doi: 10.1016/00370738(89)90015-8. Monteros, J.A., Moya, M.C., and Cuerda, A., 1993, Graptolitos AshgillianoLlandoverianos en la base de la formacion Lipeon, Sierra de Zapla, Jujuy: Su importancia en la correlation con el Silúrico de la Precordillera argentina: XII Congreso Geológico de Argentina y II Congreso de Exploracion de Hidrocarburos, Actas II, p. 304–314. Peralta, S.H., and Carter, C.H., 1999, Don Braulio Formation (late Ashgillian–early Llandoverian), San Juan Precodillera, Argentina: Stratigraphic remarks and paleoenvironmental significance: Acta Universitatis Carolinae–Geologica, v. 43, no. 1– 2, p. 225–228. Perez Guarachi, H., 1989, Contribución al estudio geológico de la formacion Cancañiri: VIII Congreso Geológico Boliviano, La Paz, Memorias, v. 1, p. 117–130. Pope, M., and Read, J.F., 1998, Ordovician metre-scale cycles: Implications for climate and eustatic fluctuation in the central Appalachians during a global greenhouse, non-glacial to glacial transition: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 138, p. 27–42, doi: 10.1016/S00310182(97)00130-2. Rapela, C.W., Pankhurst, R.J., Casquet, C., Baldo, E., Saavedra, J., Galindo, C., and Fanning, C.M., 1998, The Pampean Orogeny of the southern protoAndes: Cambrian continental collision in the Sierras de Córdoba, in Pankhurst, R.J., and Rapela, W., eds., The proto-Andean margin of Gondwana: Geological Society of London Special Publication 142, p. 181–218. Reutter, K.J., Döbel, R., Bogdanic, T., and Kley, J., 1994, Geological map of the Central Andes between 20° and 26°S, in Reutter, K.J., et al., eds., Tectonics of the Southern Central Andes: Berlin, Springer, 333 p. Rocha-Campos, A.C., 1981, Early Palaeozoic Iapó Formation of Paraná, Brazil, in Hambrey, M.J., and Harland, W.B., eds., Earth´s pre-Pleistocene glacial record: Cambridge, England, Cambridge University Press, p. 908–909. Rodrigo, L.A., Castaños, A., and Carrasco, R., 1977, La formación Cancañiri, sedimentologia y paleogeografia: Revista de Geociencias, v. 1, p. 1–22. Rust, I.C., 1981, Early Paleozoic Pakhuis Tillite, South Africa, in Hambrey, M.J., and Harland, W.B., eds., Earth´s pre-Pleistocene glacial record: Cambridge, England, Cambridge University Press, p. 113–117. Saavedra, J., Toselli, A., Rossi, J., Pellitero, E., and Durand, F., 1998, The Early Paleozoic magmatic record of the Famatina System: A review, in Pankhurst, R.J., and Rapela, W., eds., The proto-Andean margin of Gondwana: Geological Society of London Special Publication 142, p. 283–295. Saes, G.S., and Leite, J.A.D., 1993, Evolucão tectono-sedimentar do grupo Aguapeí, Proterozóico médio na porcão meridional do Cráton Amazônico:
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Schönian and Egenhoff
Mato Grosso e oriente boliviano: Revista Brasileira de Geociencias, v. 2, no. 1, p. 31–37. Saltzman, M.R., and Young, S.A., 2005, Long-lived glaciation in the Late Ordovician? Isotopic and sequence-stratigraphic evidence from western Laurentia: Geology, v. 33, no. 2, p. 109–112, doi: 10.1130/G21219.1. Sanchez, T.M., Benedetto, J.L., and Brussa, E.D., 1991, Late Ordovician stratigraphy, paleoecology and sea level changes in the Argentine Precordillera, in Barnes, C.R., and Williams, S.H., eds., Advances in Ordovician geology: Ottawa, Ontario, Geological Survey of Canada Paper 90–9, p. 245–248. Schlagintweit, O., 1943, La posición estratigrafica del Yacimiento de hierro de Zapla y la difusión del Horizonte Glacial de Zapla en la Argentina y en Bolivia: Revista Minera, Buenos Aires, v. 13, no. 4, p. 115–127. Schönian, F., 2000, Die Ablagerungsgeschichte der Cancañiri-Diamiktite (Ordovizium/Silur) im Gebiet der Santa Cruz Loma, bei Tarija, Südbolivien [diploma thesis]: Berlin, Humboldt-Universität, 139 p. Schönian, F., 2003, Ambiente sedimentario de las diamictitas de la Formación Cancañiri en el área de Sella, sur de Bolivia: Revista Técnica de Yacimientos Petroliferos Fiscales Bolivianos, v. 21, p. 131–146. Schönian, F., Egenhoff, S.O., Marcinek, J., and Erdtmann, B.-D., 1999, Glaciation at the Ordovician-Silurian boundary in southern Bolivia: Acta Universitatis Carolinae–Geologica, v. 43, nos. 1–2, p. 175–178. Schönian, F., Egenhoff, S., and Erdtmann, B.-D., 2000, Refined lithofacies codes for poorly sorted sedimentary rocks (diamictites) with an example from the upper Ordovician of southern Bolivia: International Association of Sedimentologists, 20th Regional Meeting (abstracts), p. 131–132. Schwartz, J.J., and Gromet, L.P., 2004, Provenance of a late Proterozoic–early Cambrian basin, Sierras de Córdoba, Argentina: Precambrian Research, v. 129, p. 1–21, doi: 10.1016/j.precamres.2003.08.011. Scotese, C.R., Boucot, A.J., and McKerrow, W.S., 1999, Gondwanan palaeogeography and paleoclimatology: Journal of African Earth Sciences, v. 28, no. 1, p. 99–114, doi: 10.1016/S0899-5362(98)00084-0. Sempere, T., 1995, Phanerozoic evolution of Bolivia and adjacent regions, in Tankard, A.J., Suarez-Soruco, R., and Welsink, H., eds., Petroleum basins of South America: Tulsa, Oklahoma, American Association of Petroleum Geologists Memoir 62, p. 207–230.
Sepkoski, J.J., 1996, Patterns of Phanerozoic extinction: A perspective from global data bases, in Walliser, O.H., ed., Global events and event stratigraphy in the Phanerozoic: Berlin, Springer, p. 35–51. Sheehan, P.M., 2001, The Late Ordovician mass extinction: Annual Review of Earth and Planetary Sciences, v. 29, p. 331–364, doi: 10.1146/annurev. earth.29.1.331. Sial, A.N., Dall’Agnol, R., Ferreira, V.P., Nardi, L.V.S., Pimentel, M.M., and Wiedemann, C.M., 1999, Precambrian granitic magmatism in Brazil: Episodes, v. 22, no. 3, p. 191–198. Suárez Soruco, R., 1977, Bosquejo de la Estratigrafia y Paleogeografia de la formacion Cancañiri (Silúrico) en Bolivia: Informe Interno de Yacimientos Petroliferos Fiscales Bolivianos (GXG-2156), Santa Cruz, Bolivia. Suárez Soruco, R., 1995, Comentarios sobre la edad de la formacion Cancañiri: Revista Técnica de Yacimientos Petroliferos Fiscales Bolivianos, v. 16, no. 1–2, p. 51–54. Suárez Soruco, R., and Benedetto, J.L., 1996, Primeros datos sobre la fauna de braquiópodos de la formacion San Benito en la Cordillera del Tunari, Cochabamba, Bolivia: XII Congreso Geológico de Bolivia, Tarija, Memorias, p. 211–216. Sutcliffe, O.E., Dowdeswell, J.A., Whittington, R.J., Theron, J.N., and Craig, J., 2000, Calibrating the Late Ordovician glaciation and mass extinction by the eccentricity cycles of Earth’s orbit: Geology, v. 28, p. 967–970, doi: 10.1130/0091-7613(2000)028<0967:CTLOGA>2.3.CO;2. Tassinari, C.C.G., and Macambira, M.J.B., 1999, Geochronological provinces of the Amazonian Craton: Episodes, v. 22, no. 3, p. 174–181. Toro, M., 1994, Los trilobites hirnantianos de la formación Cancañiri: XI Congreso Geológico de Bolivia, Santa Cruz, Memorias, p. 260–264. Toro, M., and Salguero, R., 1996, Los Graptolitos de la seccion Cuspidal de la formacion Tokochi, Ashgilliano temprano, en la region de Lampaya, Departamento de Cochabamba: XII Congreso Geológico de Bolivia, Tarija, Memorias, p. 589–593. Turner, J.C.M., and Mendez, V., 1975, Geología del Sector Oriental de Sta. Victoria e Iruya, Prov. de Salta, Rep. Argentina: Boletín de la Academia Nacional de las Ciencias, v. 51, no. 1–2, p. 11–24. MANUSCRIPT ACCEPTED BY THE SOCIETY 3 OCTOBER 2006
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Geological Society of America Special Paper 423 2007
Sedimentary basins in the southwestern Siberian craton: Late Neoproterozoic–Early Cambrian rifting and collisional events Julius Konstantinovich Sovetov* Anna Evgen’evna Kulikova Maxim Nikolaevich Medvedev Institute of Geology and Mineralogy, SB RAS, 3, Av. Koptyug, Novosibirsk, 630090, Russia ABSTRACT We discuss the Late Neoproterozoic and Vendian and Cambrian evolution of sedimentary basins and regional deposition events in the southwestern and southern margins of the Siberian craton from the Igarka region in the lower reaches of the Yenisei to the Patom Upland northeast of Lake Baikal. The deposition history between ca. 750 and 543 Ma was reconstructed using correlation of Late Neoproterozoic sediments with regard to diamictite deposited during global Sturtian and Varangerian glacials, rifting and collisional events, and eustatic sea level fluctuations. The accretionary and collisional events of 880–750 Ma were followed by the breakup of the supercontinent that included Siberia and the inception of failed rifts in the territory of the Yenisei Ridge, the Sayan and Baikal regions, and the Patom Upland. The main rifting stage was associated with the formation of the passive continental margin and high-energy tidal deposition on widespread broad shelves. The earliest Vendian (Ediacaran) was marked by the onset of another accretionary stage when marine deposition was disturbed by regional-scale continental glaciation (the Varangerian glacial epoch) and deep exaration of Late Riphean shelves. Vendian climate and tectonic events are reflected in the evolution of a peripheral foreland basin on the craton margin that went through several stages: (1) Early–Late Vendian subduction, glaciation, and then terrigenous carbonate shelf deposition (ca. 600 Ma). Subduction was accompanied by transgression of the fringing seas onto the craton and centrifugal transport of clastics. (2) Late Vendian collision and accretion, growth of bordering orogens, and continental molasse deposition (ca. 550 Ma). The collisional processes interrupted the marine deposition of the previous stage, closed the remnant marginal basins, and produced a peripheral orogen of a great extent. Deposition was on broad alluvial plains, and clastics were transported centripetally. (3) Late Vendian–Early Cambrian rifting and subsidence of a hangingwall clastic basin (ca. 543 Ma). Plate reorganization caused extension and related rifting and doming on the craton. The hangingwall basin on the slopes of domes received clastics from new sources on the uplifted basement blocks on the craton periphery. The dispersal of the collision-accretionary collage (the Vendian Pannotia continent) was completed in the Early–Middle Cambrian by the formation of carbonate platforms and an evaporite basin. Keywords: Late Neoproterozoic, rifigint, collision, Siberia *E-mail:
[email protected]. Sovetov, J.K., Kulikova, A.E., and Medvedev, M.N., 2007, Sedimentary basins in the southwestern Siberian craton: Late Neoproterozoic–Early Cambrian rifting and collisional events, 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. 549–578, doi: 10.1130/2007.2423(28). For permission to copy, contact
[email protected]. ©2007 Geological Society of America. All rights reserved.
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INTRODUCTION The Late Neoproterozoic section of the western and southwestern Siberian craton, in the regions of Igarka, the Sayan range of mountains, Lake Baikal, the area along the Yenisei Ridge, and the Patom Upland, records two major tectonic stages divided by a regionally extensive stratigraphic gap. The early stage encompasses Late Riphean sediments deposited under extension, and the late stage corresponds to Vendian compression and formation of a peripheral foreland basin. These stages have been called miogeocline and orogenic (Sovetov, 1977) or rifting and foredeep (Sovetov, 1993, 1997, 2001a,b) stages. Vendian subsidence and deposition in the foredeep or peripheral foreland basin have been explicitly documented in terms of lithofacies patterns and depositional systems in the Taseeva Group in the southeastern Yenisei Ridge (Sovetov, 2002c; Sovetov and Blagovidov, 2004). Early Vendian (Early Varangerian) glaciation in the Siberian craton was discussed by Sovetov and Komlev (2005). In this study we investigate the regional stratigraphy, stacking patterns, and lithology of Late Neoproterozoic (latest Riphean and Vendian) and transitional Vendian–Cambrian sedimentary systems of Siberia (foreland and evaporite basin fill) and the relationship between craton subsidence and tectonothermal events in the surrounding fold-thrust belts (Paleoasian Ocean) (Fig. 1). The critical points in the deposition history are interpreted as responses to plate reorganization associated with the assembly of the Pannotia and Gondwanaland supercontinents (Dalziel, 1997; Puchkov, 2001; Veevers, 2004). GEOLOGICAL BACKGROUND Igarka Region Late Neoproterozoic deposits in the Igarka region crop out in an anticlinal fold east of Igarka town on the Yenisei right bank (Rzhevsky and Chekhovich, 1978) (Fig. 2). The section exposed in the fold core includes the Gubinskaya Formation, of conglomerate-bearing orthoquartzite (~800 m); the Ludovskaya Formation, of phyllite and shale interlayered with metasandstone and limestone (~500 m); and an embedded block of the Igarka Formation, of pyroclastic, tuff breccia, basalt, and dolerite found also as fragments in the overlying Vendian molasse. The two formations, tentatively dated as Late Riphean (Early Neoproterozoic) from implicit evidence, are unconformably overlain by the Late Neoproterozoic Chernaya Formation (500–1000 m) and the Vendian Izluchina Formation (700–1300 m). The Chernaya Formation is composed of dolomite, limestone, and black shale intercalated with feldsparquartzose sandstone. The stratigraphic gap at its base separates the metamorphosed and sheared Early Neoproterozoic deposits from weakly altered carbonate and is marked by a dolerite sill. The Izluchina Formation, of red sandstone with siltstone at the bottom and conglomerate with coarse sandstone at the top, consists of five members (Fig. 3). The lower four members are redbeds of (1) siltstone (45–230 m), (2) siltstone and sandstone
(50–250 m), (3) sandstone (70–310 m), and (4) sandstone and conglomerate (50–500 m); member 5 is gray sandstone with mud cracks (Rzhevsky and Chekhovich, 1978). Member 4 thickens up to 500 m westward to the Basic fault, where it is distinguished as the Graviika Formation, lying unconformably on the lower Izluchina section (Shishkin, 1975). The Graviika Formation grades upward into the Sukharikha Formation, of dolomite and limestone with sandstone and marl sets (420–480 m), and into the Krasny Porog Formation, of Archaeocyatha-bearing limestone. As the basal unit of the Sukharikha Formation (and the Nemakit-Daldyn horizon), the Graviika Formation was placed at the base of the Cambrian section (Shishkin, 1975), the position that currently fits the international time scale (Gradstein et al., 2004). According to paleontological constraints and the Russian stratigraphic code, the Sukharikha Formation occurs within the Late Vendian Nemakit-Daldyn horizon, and the Krasny Porog Formation is equated to the Lower Cambrian Sunnaga horizon of the Aldan region (Luchinina et al., 1997). The Izluchina Formation is typical molasse composed of upward-coarsening fluvial deposits. The thickening of the Graviika Formation near the Basic fault and its discordant contact with the underlying Izluchina Formation is evidence of the preNemakit-Daldyn rifting in the Igarka region. Sukharikha time started with a rapid transgression recorded in a limestone section on conglomerate in the western Igarka region. Northwestern Yenisei Ridge The Vorogovka Group, with a maximum thickness of ~5400 m, fills the Vorogovka basin (southwestern Yenisei Ridge), which belongs to the system of Late Riphean failed rifts (aulacogens) coeval to the dispersal of Rodinia (Sovetov, 1993, 1997) (Fig. 4). Late Neoproterozoic rifting is inferred from the stacking pattern of sediments, the bimodal transport of clastics, initial and recurrent stages of rapid subsidence, and the connection of the rift basins to the open ocean (Sovetov, 2001b). The Vorogovka rift basin (150 × 50 km), along with the Teya-Chapa and Oslyanka basins, makes up a system of independent rifts in the Yenisei Ridge that evolved on Riphean strata of the Siberian passive margin and ophiolites thrust over the craton in the Isakovka region. The youngest ophiolites coeval to the Late Riphean rifting on the periphery of Siberia are dated 850–740 Ma (Vernikovsky et al., 2003). Vorogovka deposition was preceded by denudation of the consolidated fold-thrust belt of the Yenisei Ridge, and the Vorogovka Group is bounded from above by a regionally extensive gap. The Vorogovka basin was weakly deformed in pre-Vendian time and is overstepped, with a stratigraphic gap, by the Vendian Chapa Group. The Vorogovka Group includes four units: the Severnaya Formation (500–1500 m), with a lower and an upper subformation; the Mutnaya Formation (900 m); and the Sukhaya Formation (2400 m). The sedimentary sequence records four evolution stages of the Vorogovka failed rift (Sovetov, 2001b). Stage 1 (the Severnaya deposition) produced a graben with alluvial fans along
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Figure 1: Late Neoproterozoic geological framework of the western and southwestern Siberian craton. r.—River.
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unconformity Upper Riphean Gubinskaya Fm. Ludovskaya Fm. Igarka Fm.
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Figure 2: Generalized geology of the Igarka region, modified after (Rzhevsky and Chekhovich, 1978). Fm.—Formation; r.—River.
the sides and an alluvial delta plain in the center. The lithofacies of prodelta, delta platform, river mouth sand bars, and fluvial channels make up 10–20 m upward-coarsening cycles. Eustatic sea level rise, possibly related to spreading and transgression, produced a proto-gulf in Late Severnaya time at stage 2. Stage 3, of recurrent rifting and deposition in a deepwater basin (Mutnaya time), was associated with activity of the northeastern and then the southwestern border faults and turbidite deposition of deepsea ramp fans. Recurrent rifting can be generally related to incipient dispersal of Rodinia and lithospheric cooling; basin filling accompanied the uplift of mid-ocean ridges and faulting along the rift borders. Stage 4 yielded a broad continental embankment and a passive continental margin with tide-controlled terrigenous carbonate shelf deposition. Sediment transport was from the
uplifted rift borders and stromatolite carbonate buildups. Rapid subsidence is indicated by syndepositional deformation (seismites) and normal faults. The Vorogovka Group is sandwiched between the folded Meso- and Neoproterozoic basement west and northeast of the Vorogovka rift and the Vendian Chapa molasse fill of a broader basin, bottomed by a weak angular unconformity. The Chapa Group consists of the ~500 m Pod’yem Formation and the 1200 m Nemchanka Formation. The latter is overlain, along an erosional contact, by the Cambrian Lebyazh’ya Formation. The principal stratigraphic relations of these units were reported by Semikhatov (1962). The Pod’yem Formation includes a 300–350 m dolomite member and a 200–350 m siltstone member with chert, dolomite, and rhodochrosite interbeds (Karpinsky, personal commun.,
Vendian
Paleozoic
Neoproterozoic
Nemakit-Daldyn h.
Cambrian
Graviika Fm.
Sukharikha Fm.
I z l u c h i n a Fm.
Cambrian
Ediacaran
Cryogenian
Ludovskaya Fm.
200 m
Vor
Isakova complex
Chingasan Group
2
Chapa Group
3
67 110°
8
Siberian Craton 60°
Chistyakov Fm.
Khaerino Fm.
Ay a n k a n F m . Bol’shoi Lug Fm. Borakun Fm. Chencha Fm. Kalancha Fm. Khuzhir Fm. Kulekin Fm.
Fine-grained shelf deposits
Sandy tempestites
Sand waves and dunes
Alluvial plane with ephemeral channels Overbank fine-grained fluvial deposits
Channel deposits of deep perennial sandy braided river (downstream accretion)
Channel deposits of deep perennial gravelly-sandy braided river
Channel deposits of shallow braided and anastomosed river
Taraka massif
Katal’chikov M.
Ust’-Tagul Fm.
4
Mr
Ud
M Ma riinka Fm. Ni k - Ni kol ’sky Fm. Shan - Shangulezh Fm. Sham - Shamanka Fm. TbTirbess M. Ul Uluntui Fm. U d - U d a F m .
Calcarenite and calkrudite of mass-flow
Claye-silty limestones of storm-tidal current
Lower Proterozoic and Archean granites (basement) Sedimentary fill of Neoproterozoic failed rifts Meso-Neoproterozoic metasedimentary and metavolcanic rocks Neoproterozoic metavolcanic rocks
Base of Vendian Base of Cryogenian
Diamictites (tillites) Unconformities
Sandy turbidites
Sandy dolomites with sandstone lenses
Stromatolithic dolomites
Sandy turbidites Dolomites
Dolomite breccia
Black shales (distal turbidites)
Limestone, dolomite, dolomite breccia
Ayan
Gol
Ul
Primorie complex
Kul
Nik
Ch
Zherba Fm.
Tb
Tinnaya Fm.
7
Baikal Region Goloustnaya r., Kurtun r.
Kurtun Fm. Ushakovka Fm.
6
Buguldeika M.
Sharyzhalgai complex Muksut M.
Limestones
Chert (tuffs ?)
Sukhoi Pit and Potoskuyi Groups
Khuzh
Bol
Sham
Shan
Irkut Region, Irkut r. (3)
Ulyakha M.
5
Sayan Region Birusa r., Uda r.
(1) - modified after Rzhevsky and Chekhovich (1978), (2) - by Sovetov et al. (1975) with addition from R.B. Karpinskaya, N.I. Karpinsky and V.V. Ustalov (personal communication), (3) - by Sovetov et al. (1975) with additions from Shenfil (1991).
Ay a n Bol Bor Ch Kal Khuzh Kul -
Ostrovnoi Fm.
Veselaya Fm.
Key to symbols:
Teya and Sukhoi Pit Groups
3
Southern Yenisei Ridge Taseeva r., Angara r.
Figure 3: Correlation of Vendian molasse and its deepwater equivalent in the western and southwestern Siberian craton. Fm.—Formation; r.—River.
100°
2 4 5
1
70°
Late Neoproterozoic rifting Varanger glaciation Post-Varanger transgressions Late Vendian (Cadomian) Orogeny Early Cambrian Rifting Early Cambrian Evaporite Basin
Regional Lithostratigraphic markers Onset of:
?
?
?
1
Northeastern Yenisei Ridge Chapa r., Teya r.
Oselok Group
ISC, 2004 RusSC, 2000
Late Riphean
Northwestern Yenisei Ridge Kutukas r., Porozhnaya r. (2)
Taseeva Group
Lebyazh’ya Fm. l
Chernaya Fm.
Uglovoi Fm. T a y o z h n y Fm. Podjemskaya f.
Olkha Fm.
Redkolesnaya Fm. Greben’ Fm. A l y e s h i n s k y Fm.
Ikei Fm. Karagassy Group
Group K a c h e r g a t Fm. Baikal
Nemchanka Fm. Pod’yem Fm.
Ba ik a
La ke
Bor
Valukhta Fm. M
Dzhemkukan Fm.
Zhuya Group D a l’ n y a y a T a i g a G r o u p
Ballaganakh Group
Igarka Region Chernaya r., Sukharikha r. (1)
Kal
8
Patom Upland, Lena zone
Sedimentary basins in the southwestern Siberian craton
553
554
Sovetov et al.
96°
70°
60°
S
100°
La
ke
Ba
ik
al
Siberian Craton
N
110°
Nemchanka Group Pod’yem Fm.
Late Riphean Vorogovka Group
i r.
unconformity
e Yenis
Late Neoproterozoic
Key to symbols: Nemakit-Daldyn and Cambrian Lebyazh’ya Group disconformity Vendian
Sukhaya Fm. Mutnaya Fm. Severnaya Fm. basal layers unconformity Middle-Late Riphean Isakovka complex Metasedimentary, metavolcanic, and igneous rocks of ophiolitic association 25 km Main fault
60° Figure 4: Generalized geology of the northwestern Yenisei Ridge. Fm.—Formation; r.—River.
Sedimentary basins in the southwestern Siberian craton June 2005). The dolomite member is composed of shallow marine peritidal and supratidal facies with microphytolites, platy stromatolites, mud cracks, and tepee structures. The siltstone member was deposited in a shoreface and transitional shelf environment. The Nemchanka clastic redbeds include a ~500 m siltstone member intercalated with sandstone, a ~400 m sandstone member with siltstone layers in the upper part, and a ~300 m coarse sandstone member. Generally, the Nemchanka sequence makes up a large upward-coarsening cycle of continental (fluvial) deposition and resembles in its stacking pattern the Izluchina Formation of the Igarka region. Thick channel fill of braided fluvial systems at the base of members 2 and 3 records the abrupt cratonward progradation of rivers that flowed northward, as is indicted by measurements of current directions. The similarity of the Nemchanka and Izluchina Formations is stressed by the change from sand to sandygravelly river systems in Late Izluchina–Early Lebyazh’ya time. A ~300-m thick member of sand-gravel fluvial deposits at base of the Lebyazh’ya Formation grades into a member of intermittent shallow marine microphytolithic dolomite, sandy dolomite, and sandstone, i.e., the stacking pattern is identical to those of the Graviika and Sukharikha Formations from the Igarka region. Northeastern Yenisei Ridge Late Neoproterozoic deposits in the northeastern Yenisei Ridge overlie, with a prominent stratigraphic unconformity, the metamorphosed Meso- and Neoproterozoic rocks of the Sukhoi Pit and Tunguska Groups (Semikhatov, 1962), intruded by ca. 850 Ma granites (Nozhkin et al., 1999) (Fig. 5). The Late Neoproterozoic section encompasses the Late Riphean Chingasan Group (2000 m) and the Vendian Chapa Group (2000 m), which fill the Teya-Chapa failed rift (Sovetov, 1997), and the overriding East Yenisei foredeep (Sovetov, 2002a), respectively. Granites dated 750–720 Ma intrude the basement but not the Chingasan Group (Vernikovsky et al., 2003), and the latter thus should be younger than 720 Ma. The Chingasan Group includes the Lopatinsky (450–1100 m), Kar’yer (450–1000 m), and Chivida (500– 800 m) Formations. Fluvial and deltaic deposition of the Lopatinsky Formation gave way to rapid ingression. A steadily subsiding shelf that developed in Kar’yer time under a strong tidal influence was overdeepened by a rifting pulse that also produced local alluvial fans. The Chivida recurrent rifting formed a deep basin and deep-sea fans and was associated with bimodal volcanism in the northwestern Teya-Chapa rift and a glaciation equated to the Sturtian ice age. The subsidence and ingression stages are similar to those of the Teya-Chapa and Vorogovka rifts, while the shelf and recurrent rifting stages initiated and ended independently in each basin. The clastic redbeds of the Yenisei Ridge and the underlying Pod’yem dolomite were identified in the 1960s and 1970s as the Nemchanka Formation (Semikhatov, 1962; Khomentovsky et al., 1972; Sovetov, 1977). Later the Chapa Group was distinguished within the same stratigraphic range (Butakov et al., 1975), and it is currently acknowledged as a valid regional stratigraphic unit
555
for the northern Yenisei Ridge. The Chapa Group in this region consists of the Suvorovsky, Pod’yem, and Nemchanka Formations (listed upsection). The Suvorovsky Formation is composed of outsized red-colored feldspar-quartzose sandstone with thin interlayers of fine dolomite tempestite (350–500 m) and discordantly overlies the Chivida Formation, with the contact marked by boulder conglomerate lenses (Khomichev et al., 2002). The Suvorovsky formation pinches out outside the Teya-Chapa basin and leaves the Pod’yem Formation to rest on metamorphosed Meso- and Neoproterozoic exposures on the rift borders. The Pod’yem Formation is composed of heterolithic terrigenous carbonate shallow marine deposits (200–320 m) and grades upsection into the red sandstone (1600–4200 m) of the Tayozhnaya and Uglovoi Formations (Sovetov, 1977). The Pod’yem Formation is richer in sandstone than its equivalents of the same name in the West Yenisei foredeep, which indicates an intracratonic provenance of clastics. It is bottomed by a member of mixtite or breccias that we suppose are equivalent to the tillites found in the Sayan and Baikal foredeeps (Sovetov, 2002c). The Pod’yem Formation is gradually transient into the Tayozhnaya Formation, divided into three members making up an upward-coarsening sequence. The lowermost member (400–500 m) is dominated by siltstone and fine sandstone with scarce coarse sandstone lenses; the intermediate member (500–600 m) is composed of typical 1– 5 m sand cycles deposited by a braided river; the upper member (300 m) is a succession of 2–3 m sandstone cycles deposited by a river with alternate bars and contains minor amounts of fine clastic components. Numerous measurements of current directions showed fans oriented mostly to the NNE or the SSE (Sovetov, 2002c; Sovetov and Blagovidov, 2004). The overstepping red-colored Uglovoi Formation was distinguished within the Nemchanka Formation (Sovetov, 1977), proceeding from the presence of a conglomerate (10–75 m) and coarse channel sand (200 m) member grading upward into a 50 m member of shallow marine gray sandstone with glauconite and biothermal stromatolite dolomite below a 300 m redbed of shallow marine sandwave and dune facies. Upsection there follow dolomite and dolomite breccia of the Lebyazh’ya Formation, which was assigned an Early Cambrian age by correlation with the salt- and anhydrite-bearing Usol’ye Formation. The stratigraphy of the Tayozhnaya and Uglovoi Formations is similar to those of the Nemchanka Formation in the West Yenisei foredeep and the Izluchina Formation in the Igarka foredeep. The section exhibits three stages (Middle Tayozhnaya, Late Tayozhnaya, and Early Uglovoi) of alluvium progradation and tectonic activity of the surrounding orogen. The K-Ar glauconite ages of Late Neoproterozoic sediments obtained before the 1990s are ambiguous and have not been considered here. Southern Yenisei Ridge The Vendian section of the region is represented by the Taseeva Group lying over the deformed Meso- and Neoproterozoic
556
Sovetov et al.
70°
Siberian Craton
95°
100°
La
ke
Ba
ik
al
60°
110°
60
Key to symbols: Cambrian Lebyazh’ya Fm.
ar . ap
Pod’yem Fm.
Ch
Chapa Group
Tayozhny and Uglovoi Fm. Suvorovsky Fm. disconformity
ar .
Late Riphean Trachy-basalt and tuf f Chivida. Fm. Chivida Fm. Kar’yer Fm.
Te y
Chingasan Group
Late Neoproterozoic
Vendian
Lopatinsky Fm. unconformity Middle-Neoproterozoic metasedimentary and igneous rocks
25 km
Main faults Figure 5: Generalized geology of the northeastern Yenisei Ridge. Fm.—Formation; r.—River.
Sedimentary basins in the southwestern Siberian craton strata of the Siberian craton margin and Lower Proterozoic granite of the Tarak Complex. The structural position of the Taseeva Group generally shows no inheritance of Vendian foredeeps from the Late Neoproterozoic failed rifts, but a fragment of the IyaTumanshet rift represented by the Khaerino feldspar-quartzose gravelstone, sandstone, and siltstone over the Tarak granite is found in the southern Yenisei Ridge (Fig. 6). The Taseeva rocks are Late Precambrian molasse of both interior and exterior provenance. The upper Taseeva section was deposited in a syncollisional basin when orogens appeared along the entire southwestern periphery of the Siberian craton and river-transported clastics filled the South Yenisei foredeep. The scrutinized depositional systems of the Taseeva Group bear evidence of current directions and cyclic sequences of different scales (Sovetov and Blagovidov, 2004). The Taseeva Group includes five regional units, namely, the Aleshinsky (930 m), Chistyakov (270 m), Greben’ (413 m), Veselaya (233 m), and Redkolesnaya (295 m) Formations. The Aleshinsky Formation of red fluvial deposits consists of six members: al11 (140 m), al12 (240 m), al13 (200 m), al21 (70 m), al22 (180 m), and al23 (70 m), deposited by river systems of different types or different deposition regimes (Sovetov, 2002b; Sovetov and Blagovidov, 2004): (1) a gravelly braided system (140 m al11) that developed in a single valley or in several nested valleys, with boulder conglomerate at its base as found in the northern South Yenisei foredeep (Khomichev et al., 2002); (2) a sandy meandering system of channels with lateral accretion bars and interchannel fines (240 m); (3) uncertainly bedded unsorted siltstone and sandstone of a distal fluvial plain (200 m) and massive sets of siltstone in association with ephemeral sheetflood deposits void of flow textures, looking like loess deposited in an arid climate (this member correlates with glacial deposits discovered at the same stratigraphic level in the Sayan region [Sovetov, 2002b]); (4) a gravelly anastomosing system of gravel and sand deposited in single or stacked (two or three) channels and interchannel siltstone sets (25–30 m); (5) a sandy ephemeral channel system mostly of fine sheetflood deposits and inferior sand fill of small shallow (10–50 cm) ephemeral streams (40 m); (6) a gravelly-sandy meandering system (100 m) with overbank deposits and crevasses in its top (80 m); (7) a sandy braided system (26 m) of deposition cycles with weakly developed overbank fines (stacked channel structure); and (8) a delta system strongly different from the previous ones in absolute predominance of fine deposits (47 m). The Aleshinsky fluvial deposition was associated with a fourfold increase in erosion rate and river dynamics. The progradation of alluvium, as well as the coarse grain sizes of the channel fill, may have been caused by activity of the source areas and reactivation of the drainage network during deglaciation simultaneous with the rise of the erosion base level as the effect of a sea level change. The transport of clastics was in the northwest-northeast or southeast-south directions (Sovetov and Blagovidov, 2004). The Chistyakov Formation, of gray shallow marine clastics, includes three members: (1) cst1 (93 m), black and dark gray siltstone and limestone with scarce coarse cross-bedded sandstone interbeds and sporadic clayey dolomite layers; (2) cst2 (99 m),
557
dark gray siltstone and wavy-bedded sandstone of storm structures (bottom) and white and light gray massive coarse and fine cross-bedded sandstone with oblique sandwave foreset bedding (top); (3) cst3 (78 m), petricolored outsized sandstone with storm structures (bottom) and thinly alternated red siltstone and sandstone with lenses of gray fluvial sandstone densely cut by mud cracks (top). The members constitute upward-coarsening thirdorder cycles (Sovetov and Blagovidov, 2004) related to eustatic sea level changes that are likewise recorded in the Pod’yem Formation in the western and eastern Yenisei Ridge and in the Marnya and Uda Formations of the Sayan region. The Greben’ Formation is composed of fluvial deposits, mostly mixed interchannel sandstone and siltstone. It includes four cyclic members: gr1 (90 m), gr2 (68 m), gr3 (106 m), and gr4 (93 m), each bottomed by a thick layer of massive cross-bedded sandstone with siltstone intraclasts. The stacked channels in gr1 can serve as stratigraphic markers. The Veselaya Formation, of red channel fill, differs in the absence of relatively thick siltstone layers; alternating bars of fine- to coarse-grained fluvial sandstone build stacked successions of channels. Scarce siltstone sets separate members vs1 (97 m), vs2 (59 m), vs3 (55 m), and vs4 (49 m). The Redkolesnaya Formation, of shallow marine and shoreface bar and dune facies, is divided into seven members, each deposited in a cycle: rl11 (80 m), rl12 (58 m), rl13 (32 m), rl21 (35 m), rl22 (27 m), rl23 (28 m), and rl24 (25 m). The members are coarser at the bottom, and their upper sections are composed of sandstone with graded bedding and lamination. Conglomerate lenses and abundant sand attest to the proximity of alluvial fans. Redkolesnaya deposition occurred mainly under orogenic pulses in the conditions of high sea stand and shore retreat associated with sand expansion over the beach. Sediments were deposited in a peritidal environment on the upper beach, at different levels of the shoreface zone, then transitioned to the inner shelf, under high-wave activity and relatively weak tides, and eventually produced a broad strand plain (Sovetov and Blagovidov, 2004). The Redkolesnaya Formation is a lithologic and stratigraphic equivalent of the lower Uglovoi Formation, the basal member of the Lebyazh’ya Formation in the northern Yenisei Ridge and of the Graviika Formation in the Igarka region. It represents the start of a new deposition stage in the foreland basin evolution and makes up a large cycle together with the overlying terrigenous carbonate Ostrovnoi Formation. Its stratigraphic equivalents are found in the northern and eastern Siberian craton (the Staraya and Manykai Formations and the Ust’-Yudoma Formation) (Khomentovsky et al., 2004), and it correlates with the upper section of the Yudoma Group of the Aldan Shield (Sovetov et al., 1975). The Ostrovnoi dolomite formation overlying the Redkolesnaya Formation contains small-shell fauna and was equated to the Nemakit-Daldyn horizon on a chemostratigraphic basis (Khomentovsky et al., 1998, 2004). The NemakitDaldyn horizon is placed at the Late Vendian in the Russian stratigraphic code (Khomentovsky and Karlova, 2005) and at the base of the Cambrian on the international time scale, and it is associated with a tectonic event that restructured the sedimentary basins at
558
Sovetov et al.
94°
70°
58° An
Siberian Craton 60°
580
Motygino gar
a r.
ik al
Q
Tas e
Ba La ke
100°
eva
110°
Key to symbols: Q
r.
Q
Quaternary unconformity
Mesozoic unconformity
25km
Early-Middle Paleozoic Late Nemakit-Daldyn and Cambrian Redkolesnaya Fm. (Early Nemakit-Daldyn) Veselaya and Greben’ Fm.
Taseeva Group
Vendian
unconformity
Chistyakov Fm. Aleshinsky Fm. unconformity
Early-Middle Neoproterozoic Mesoproterozoic Figure 6: Generalized geology of the southern Yenisei Ridge. Fm.—Formation; r.—River.
543 Ma (Brasier 1992; Bowring et al., 1993). According to sedimentological evidence, the Redkolesnaya Formation is a basal unit of a regional deposition cycle and marks the inception of a new sedimentary basin. Therefore, it appears reasonable to correlate its base to a global-scale boundary, namely, the lower boundary of the Cambrian (Zharkov and Sovetov, 1969; Sovetov, 2004). Sayan Region Neoproterozoic deposits in the northeastern foothills of the East Sayan chain of mountains are represented by the Late Neoproterozoic Karagassy Group (2400 m), which fills the Late Riphean Iya-Tumanshet failed rift, and by the Vendian Oselok Group (max to 2700 m), which fills the peri-Sayan foredeep. The
Iya-Tumanshet rift developed on metamorphosed sediments and volcanics of the Early Neoproterozoic Biryusa Group, intruded by two granite plutons dated between 1870 and 1734 Ma (Turkina et al., 2003), and on Early and Middle Proterozoic metasedimentary and metavolcanic fill of the Urik-Iya graben in the southeast (Fig. 7). The Karagassy Group rests unconformably on heterochronous basement complexes and comprises the Shangulezh (350–400 m), Tagul (450 m), and Ipsit (800 m) Formations (Bragin, 1986; Sovetov and Blagovidov, 2006). The Shangulezh Formation is composed of alluvial fan facies in grabens and dolomite deposited during the initial sea ingression. The Tagul Formation consists of shallow marine stromatolitic and microphytolitic dolomite, sandwaves, and dunes. The Ipsit Formation is made up of
Sedimentary basins in the southwestern Siberian craton
559
Key to symbols:
56°
Jurassic
70°
unconformity
Taishet
C1
Lower Carboniferous
Siberian Craton
unconformity
60°
Middle Devonian ik a
l
unconformity La ke
r.
Ba
Lower Ordovician
unconformity
110°
Tag ul
100°
Cambrian Talaya Fm.
Nizhneudinsk
Bi
ryu
sa
r.
Vendian
Nemakit-Daldyn Ust’-Tagul Fm. Oselok Group
disconformity Late Rephean
54°
Karagassy Group
ar .
Iya
r.
Lower Proterozoic Biryusa Group and Sayan igneous complex Faults
р.
Ud
98°
unconformity
Main fault
100° Figure 7: Generalized geology of the Sayan region. Fm.—Formation; r.—River.
50 km
carbonate and terrigenous tempestite of shoreface and transitional inner shelf origin. The Karagassy Group is cut by Nersa dolerite sills and intrusions correlated to mafic dikes in the Sharyzhalgai basement block; the Ar-Ar and Sm-Nd ages of gabbro-dolerite are within 740–780 Ma (Gladkochub et al., 2003). The Oselok Group includes the Marnya (400–660 m), Uda (200–550 m), and Aisa (1500 m) Formations. The Aisa Formation has been distinguished as joining with the upper Uda Formation as the Ikei Formation (Bragin, 1985) to make a single supersystem of fluvial deposition (Sovetov, 2002b; Sovetov and Blagovidov, 2004). The Marnya Formation overlies the Karagassy Group,
with a stratigraphic gap and a prominent erosional contact over the pre-Vendian exaration surface (Sovetov, 2002c; Sovetov and Komlev, 2005). Glacial deposits include diamictite (tillite), boulder breccia, boulder conglomerate, fluvial sandstone with pebbles, aeolian sandstone, and lacustrine black shale. Glacial and subglacial facies replace each other and serve as the basal member in all localities we studied. The Marnya Formation consists of eight members (Sovetov, 2002a; Sovetov and Komlev, 2005): (1) the Nersa member, of glaciofluvial massive boulder conglomerate (19 m); (2) the Ulyakha member, of diamictite (tillite) (15–55 m); (3) the Tygnei member, of lacustrine
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Sovetov et al.
black pyritized bituminous mudstone and siltstone with lenses of fluvial sandstone (80–90 m); (4) the Plity member, of tillite and fluvial and aeolian white cross-bedded quartzose sandstone (20–35 m); (5) the Kedrovy member, of glaciofluvial coarse litharenite or extralitharenite cross-bedded sandstone with conglomerate lenses or isolated well-rounded pebbles and cobbles, at the base of the fluvial cycles (40 m); (6) the Ozerki member, of laminated cap dolomite, shelf dolomite, microphytolite and stromatolite dolomite, dune and sandwave sandy dolomite, and dark gray bar orthoquartzites (15–20 m), in the middle part bearing abundant ichnofossils (max. to 110 m); (7) the Ognit member, of dark gray and white shoreface arkose sandstone and orthoquartzite with coarse sandstone lenses on the top of dunes and sandwaves (30–90 m); and (8) the Bol’shaya Aisa member, of shelf sandstone and siltstone with hummocky and horizontal bedding of tempestite bearing scarce fossils of metazoan animals (100–270 m). The depositional systems of the Marnya Formation were oriented toward the marginal sea and received clastics from the central Siberian craton. Glaciation apparently had two centers: an outer center in the accretionary belt and an inner one in the craton interior, as is indicated by the current directions. The Uda Formation includes (1) the Lower Uda member, of fluvial feldspar-quartzose sandstone and coarse sandstone grading into coastal bar sandstone in the southwest (15–70 m); (2) the Unyl member, of deltaic siltstone with sandstone lenses deposited in subsidiary channels and river mouth bars (25–100 m); (3) the Peschernaya member, of dark gray bituminous stromatolithic, microphytolitic, and silt-grained hummocky-bedded limestone (7–30 m); (4) the Kagat member, of massive cross-bedded fluvial sandstone, sets of channel bar and overbank siltstone, and sandstone of shallow crevasse channels on a delta plain (50– 70 m); and (5) the Muksut member, of massive thickly laminated cross-bedded quartz-litharenite fluvial sandstone (100–150 m). The current directions in the Lower Uda fluvial deposits and the change from alluvial to shallow marine facies indicate that the accommodation basin was located southwest of the Siberian craton. The Kagat and Muksut members are joined with the overlying Aisa Formation into a single fluvial complex. The Aisa Formation is composed of fluvial sediments deposited by perennial deep river and anastomosing systems that make up tens of cycles strongly dominated by overbank fine clastics. The uppermost Aisa section is formed by the Katal’chikov member (~200 m), which marks the activity of the source area and the appearance of a braided system style. The Aisa continental deposits, along with the Kagat and Muksut members of the Uda Formation, are syncollisional molasse produced by erosion of Late Vendian flanking orogens. There is regional correlation with two levels (the Muksut and the Katal’chikov) of prominent cratonward progradation of fluvial systems. The Muksut member is the key marker in the peri-Sayan foredeep (Bragin, 1985) and is also traceable in other foredeeps of the southwestern craton periphery (Sovetov, 2002c; Sovetov and Blagovidov, 2004).The Aisa Formation is overlain with a disconformity by the Ust’-Tagul Formation, consisting
of two subformations: a fluvial (~100 m) unit and a shallow marine and lagoonal terrigenous carbonate (~100 m) unit. The lower Ust’-Tagul subformation consists of cyclic gravelly and sandy braided-system lithofacies with minor amounts of fine clastics (Fig. 8). The fluvial system at the lower-upper subformation boundary gives way to tidal plain and lagoonal facies with abundant trace fossils of typical Early Cambrian Metazoa. The Ust’-Tagul Formation of clastics is very similar to the deposits of the Redkolesnaya, Uglovoi, Lower Lebyazh’ya, and Graviika Formations from the Yenisei Ridge and the Igarka region in its stratigraphic position, the erosive gap at its base, and the abrupt changes of its sedimentary systems. Irkut Region No deposits of Late Neoproterozoic failed rifts are known in the Irkutsk region (Fig. 9). Vendian sediments of 450–750 m, including the Olkha Formation and the Moty Group, cover the Sharyzhalgai basement uplift, a block of Archean metamorphic and igneous rocks (also called the Irkut or Kitoi uplift in Khomentovsky et al., 1972, and Sovetov, 1977). The Vendian stratigraphy of the region has been a point of discussion (see Khomentovsky et al., 1972; Sovetov, 1977). The Vendian section is thinner than the fill of foredeeps, and its different horizons overlap the Sharyzhalgai block. The Late Neoproterozoic section consists of the 75–500 m Tyret’ (Olkha) Formation and the Moty Group, divided into the Khuzhir (70–280 m), Bol’shoi Lug (70–150 m), and Shamanka (Nurtei) (160–250 m) Formations (Sovetov, 1977). The Olkha Formation is composed of shallow marine siltstone and sandstone with dolomite lenses and has a breccia or sandstone at its base, with inclusions of angular quartz and martite pebbles (up to 10 m). The Khuzhir Formation overlies the deeply eroded surface of the Olkha Formation and is likewise bottomed by a 10-m thick layer of conglomerate containing pebbles of Archean rocks; the overlying coarse cross-bedded sandstone was deposited by braided-system channels. The Bol’shoi Lug Formation is made up of white tidal fine quartzose sandstone. The Shamanka Formation comprises two members: sham1, of red fluvial coarse gravelly sandstone of an alternate bars system (100 m) (Fig. 10), and sham2, of red and yellow shallow marine fine sandstone of dunes and sandwaves (40 m) (Fig. 11). It underlies the terrigenous carbonate Irkut Formation, deposited as coastal bars and back-bar lagoons. The transgressive quartz sandstone of the Bol’shoi Lug shoreface and fluvial plane deposits of the Shamanka Formation are well known; the sandstone is widespread in the inner Siberian craton as the Parfenovka member (Sovetov, 1977; Tyschenko, 1980; Shenfil’, 1991). The second transgression after the far cratonward progradation of the Shamanka alluvium partly reworked fluvial sediments into a shoreface complex. The most progressive stage of this transgression is marked by the Irkut Formation, with its base correlated to the geophysical tiepoint M2 (Sovetov, 1977; Tyschenko, 1980; Mel’nikov, 1982; Shenfil’, 1991). The Irkut uplift separates the peri-Sayan and peri-Baikal foredeeps, which are filled with the Oselok and Baikal Groups.
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Figure 8: Generalized (A) and detailed (B–E) stratigraphy of the Ust’-Tagul Formation, Sayan region (Tagul and Biryusa rivers). Fm.—Formation.
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Figure 9: Generalized geology of the Irkut region. Fm.—Formation; r.—River.
Correlation of the Olkha Formation and the Moty Group to the Vendian fill of the two foredeeps is difficult because of facies change and a threefold thickening of the deposits. Correlation can be based on three stratigraphic ties: (1) Olkha marine, (2) Khuzhir continental (fluvial), and (3) Shamanka continental (fluvial) and shallow marine facies. The Olkha Group generally correlates with the Marnya and Uda marine deposits, the Khuzhir Formation bears many features of the Aisa fluvial deposits, and the Shamanka Formation is similar to the lower Ust’-Tagul Formation in the structure and lithology of fluvial deposits. The Ust’-Tagul Formation, in its full range, and the Shamanka Formation, together with the Irkut Formation, were jointly interpreted as a single unit with its base or top correlated to the Lower Cambrian boundary (Zharkov and Sovetov, 1969; Khomentovsky et al., 1972, 1998). The Shamanka Formation has the same stratigraphic position as the Redkolesnaya Formation in the southern Yenisei Ridge and can be attributed, in view of the previously mentioned considerations, to the lower part of the Nemakit-Daldyn biostratigraphic unit.
Baikal Region The Baikal Group was first identified in the western Baikal region and divided into the Goloustnaya, Uluntui, and Kachergat Formations (Tetyaev, 1916; Fig. 12). Later it was traced in the Primorsky Ridge as far as the northern tip of Lake Baikal (Mats, 1962), and its stratigraphic division and age estimates were only slightly refined (Khomentovsky et al., 1972; Maslov, 1983; Shenfil’, 1991). The age was estimated against microphytolite and stromatolite as Late Neoproterozoic (Baikalian) (Khomentovsky et al., 1972; Shenfil’, 1991; Dol’nik, 2000; Khomentovsky, 2002) and later refined as latest Neoproterozoic (Vendian) on the basis of sequence stratigraphy correlation with the Oselok Group of the peri-Sayan foredeep and correlation of the tillite horizon at the base of both groups with those of the Varangerian and Marinoan tillites (Sovetov and Komlev, 2005). The Baikal Group rests discordantly on Early Proterozoic granite of the Primorsky Complex, with a structural gap in the Western Baikal region; in the northeast it oversteps a 500–700-m
Sedimentary basins in the southwestern Siberian craton
Key to symbols: Medium-grained sandstone Coarse-grained sandstone Syndepositional deformation (seismites ?)
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Figure 10: Fluvial channel cross-bedded and parallel-bedded deposits of the (A, B) Shamanka Formation (Irkut River) and (C)—Ushakovka Formation (Kurtun River).
thick lens of the volcanoterrigenous Nugan Formation, which overlies the basement and pinches out to the edges of the periBaikal foredeep (Maslov, 1983). The structural position of the Nugan Formation is similar to that of the Late Neoproterozoic (pre-Vendian) sedimentary and volcanosedimentary fill of failed rifts in the Yenisei Ridge. The Baikal Group consists of the Goloustnaya (470–500 m), Uluntui (400–600 m), and Kachergat (650–2000 m) Formations. Southwest of the Irkut uplift, the Goloustnaya Formation pinches out, the Uluntui Formation thins down and gives way to the Olkha Formation (350–400 m), and the Kachergat Formation abruptly thins to 180 m and is replaced by the Khuzhir Formation. The Goloustnaya Formation includes four members: (1) the Buguldeika diamictite (tillite) (56 m), with up to 1.5 m granite blocks and boulders over Early Proterozoic granite; (2) a cap dolomite member with shale sets (60 m); (3) a sandy dolomite and quartzitic sandstone member with dune and sandwave structures (80–100 m); and (4) a microphytolitic bar dolomite member (~200 m). The terrigenous carbonate sediments above
diamictite were deposited in tidal and inner shelf environments. The Uluntui Formation is composed of shallow marine deposits and includes five cyclic members (sequences) grading from low-stand terrigenous to high-stand carbonate deposits: (1) ul1, of dark gray quartz, coarse sandstone, and orthoquartzite, with dolomite on the top (110 m); (2) ul2, of phyllite with dolomite on the top (90 m); (3) ul3, of phyllite with dolomite on the top (70 m); (4) ul4, of phyllite with dolomite on the top (100 m); and (5) ul5, of sandstone with sandwave structures and tempestite, with dolomite on the top (80 m). Phyllitic fine and coarse sandstones erosively overlie dolomite; the lower boundaries of sequences ul1 and ul5 can serve as markers. The Kachergat Formation, of fluvial, deltaic, and shelf deposits, makes up a clastic assemblage of syncollisional molasse and includes four members (Shenfil’, 1991) traceable in the peri-Baikal foredeep: (1) kc1, of coarse channel sandstone (60–150 m); (2) kc2, of deltaic siltstone and argillite with small channel-fill lenses (300–400 m); (3) kc3, of channel and mouth bar sandstones (60–100 m); and (4) kc4, of shoreface tempestite and sandwave
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Figure 11: Generalized (A) and detailed (B–E) stratigraphy of the Shamanka Formation and the lower part of the Shankhar Formation, Irkut region (Irkut River).
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Figure 12: Generalized geology of the Baikal region. Fm.—Formation; r.—River.
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(bar) sandstone, with transitional black tempestite on the top (~500 m). The Kachergat Formation records two large events (kc1 and kc3) of cratonward progradation of braided and alternate bar systems. NNE and NNW river current directions in the western Baikal region were inferred from intrachannel dune and bar foresets, i.e., the orogen was located south and southwest of today’s Lake Baikal (Sovetov, 2002c; Sovetov and Blagovidov, 2004). The Ushakovka Formation, of fluvial deposits with conglomerate at the base, erosively overlies the Kachergat Formation (Mats, 1962; Khomentovsky et al., 1972; Sovetov, 1977; Shenfil’, 1991) and demonstrates the farthest cratonward progradation of fluvial systems similar to the systems of the Shamanka, Ust’-Tagul, Uglovoi, and Graviika Formations. The Ushakovka Formation includes a ~20 m member of fluvial, occasionally boulder, conglomerate channel fill (Fig. 13) and a 200 m (over 1000 m in the central peri-Baikal foredeep) member of coarse fluvial sandstone (Shenfil’, 1991) deposited by deep and small braided rivers. The Ushakovka fluvial system grades upsection into the Kurtun Formation, of shallow marine terrigenous deposits (~200 m); the abrupt transition is similar to the transition from fluvial to shallow marine facies in the upper sections of the Shamanka, Ust’-Tagul, and Uglovoi Formations. The second transgression episode is recorded in the southwestern craton periphery and interior by widespread tidal plain and lagoonal crabonate and terrigenous carbonate sediments; the facies change is marked by the geophysical tiepoint M2 (Sovetov, 1977; Tyschenko, 1980; Mel’nikov, 1982; Shenfil’, 1991). Patom Upland The Patom Upland has a complex tectonic structure and history. Our study is restricted to the northwestern Lena zone of the upland, which bears easily detectable diagnostic lithological units that enable us to tie the strata of the Neoproterozoic passive margin of Siberia to the fill of failed rifts and foredeeps (Fig. 14). The Early Proterozoic (1950–1750 Ma) basement is overlain, along a tectonic contact, by the 150–950 m Purpol Formation, of shallow marine orthoquartzite, quartz conglomerate, and aluminous shale (Ivanov et al., 1995). The formation is assumed to make up the lower section of the Teptorgo Group and to be a fragment of an Early Riphean platform cover. It is overstepped by the Medvezh’ya Formation, identified either as the upper unit of the Teptorgo Group or as the basal unit of the overlying Ballaganakh Group. According to our field data, the Medvezh’ya Formation is impossible to tell from the overlying Khorlukhtakh Formation by their lithology of coarse clastics. Therefore, the over 10,000-m thick sedimentary sequence of the Late Proterozoic continental margin should start with graben fill and the related Medvezh’ya dolerite sills and basalts. Rifting is discussed in detail by Ivanov et al. (1995). The early rifting stage (1000–800 Ma) was completed with a collision of island arcs developed on continental crust with continental terranes (Rytsk et al., 2001), deformation, metamorphism of volcanosedimentary rocks, and plutonism. The second rifting stage (800–620 Ma) was marked by production of
juvenile crust and opening of the Paleoasian Ocean. The Patom passive margin was subsiding continuously, and continental slope and foot environments persisted as long as the Late Vendian. The Patom Neoproterozoic section consists of the 2000– 4000 m Ballaganakh, 2500–3500 m Dal’nyaya Taiga, and 750– 1200 m Zhuya Groups. The Ballaganakh Group includes the Medvezh’ya Formation, of diamictite, orthoconglomerate, dropstone-bearing shale, basalt, and tuff (150–600 m); the Khorlukhtakh Formation, of diamictite, turbidite, orthoconglomerate, and dropstone-bearing shale (300–800 m); the Khaiverga Formation, of terrigenous and carbonate turbidite and black shale, occasionally with dropstones (500–1500 m); the Bugarikta Formation, of sand turbidite and debris flow deposits with black shale interbeds (800 m); and the Mariinka Formation, of carbonate and terrigenous carbonate turbidite and olistostromes (150–650 m). The Ballaganakh Group shares many features of similarity with the fill of Neoproterozoic rifts, namely (1) boulder orthoconglomerate and diamictite (tillite), (2) carbonate shelves and carbonate mud redeposited by turbidity and mass flows, (3) mafic and felsic magmatism, (4) recurrent rifting that produced deepsea fans, and (5) carbonate platform with clastics on its slope. Dropstones and layers with inclusions of unsorted redeposited material indicate widespread long-lasting glaciation correlated with the Sturtian glacial epoch proceeding from the presence of a younger layer of glacial deposits at base the Dzhemkukan Formation (Chumakov, 1993). The Dal’nyaya Taiga Group includes the 650–1000 m Dzhemkukan Formation, of diamictite, sand and sand-mud turbidite, and black shale; the 150–400 m Barakun Formation, of carbonate turbidite and mass flow deposits; and the 1500–2000 m Valyukhta Formation, of sand-mud turbidite, prodelta black shale, sandstone of subaqueous and surface channels, and limestone of microphytolite banks. The Dzhemkukan Formation, stripped in some sections, consists of diamictite with a signature of glacial deposition (Chumakov, 1993). The lower part of the Barakun Formation looks like cap dolomite. The Valyukhta rocks form a large marine delta with its provenance in the Siberian craton. The three depositional systems correlate well with the Goloustnaya and lower Uluntui Formations, which fill the peri-Baikal foredeep. The good markers are the diamictite (tillite) at the base of the Dzemkukan and Goloustnaya Formations and the stromatolite and microphytolite carbonate platform in the upper part of the Uluntui and Valyukhta Formations. This last unit in the Patom Upland is named the Kalancha Formation (Khomentovsky et al, 1972). The Zhuya Group comprises the Nikola Formation, of siltstone and marl with tidal-current structures (350 m), and the Chencha Formation, of stromatolite biohermal frameworks, microphytolite banks with sandstone lenses, and carbonate tempestite (500–600 m). The Chencha deposits demonstrate a radical facies change toward the Baikal region, and therefore this stratigraphic level is more vague in terms of correlation. We believe that the most significant feature of the Zhuya Group deposits as a whole is the progradation of terrigenous clastics in the Siberian craton. The Kulekin Member of the Kalancha Formation, of
Sedimentary basins in the southwestern Siberian craton
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Figure 13. Fluvial channel of cross-bedded conglomerate and sandstone in the Ushakovka Formation, Baikal region (Kurtun River).
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Figure 14: Generalized geology of the northwestern Patom Upland, Lena zone. Fm.—Formation; r.—River.
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conglomerate and sandstone above limestone, would mark the beginning of this process. In the Baikal region we correlate the first member of the coarse-grained sandstone in the Kachergat Formation with Kulekin conglomerate as a result of a coeval orogeny outside the Siberian craton. The Zherba Formation (350–500 m) is composed of shallow marine quartzose sandstone, coarse sandstone with a black limestone and shale member in the middle, and a coarse and fine sandstone member at the top. Conglomerate and coarse sandstone in the upper Zherba Formation and the overlying terrigenous carbonate Tinnaya Formation make up a cyclic sequence similar to that of the Shamanka and Irkut Formations and of the Ust’-Tagul Formation and its equivalents. The Patom passive margin evolved in several stages, from grabens through a deepwater basin to a broad tidal shelf with stromatolite biotherms and biostromes. Numerous measurements reveal eastward transport of clastics from granitoid and metamorphic basement complexes from the craton into the opening Paleoasian ocean. The Ballaganakh and Dal’nyaya Taiga sections were mostly deposited in euxinic deep-sea environments. PreVendian Late Neoproterozoic rifting included an initial stage (in Medvezh’ya-Khorlukhtakh time) associated with mafic magmatism and glaciation (during the Sturtian epoch) and a recurrent rifting stage (in Bugarikta time) that produced a chain of coarse clastic fans (Sovetov et al., 1999). Glaciation and a sea level fall in Dzhemkukan time (Early Vendian) provided the conditions for voluminous redeposition of unsorted clastics onto the continental slope foot (Chumakov, 1993). This later glaciation correlates with the Varangerian glacial epoch (Sovetov and Komlev, 2005). The Late Varanger rifting (in Valyukhta time) led to the advance of large sand-pelitic deltas. The lower sea-level stages alternated with stages of broad margin subsidence, and formation of carbonate shelves (in the Late Khaiverga, Mariinka, Barakun, Late Valyukhta, and Chencha ages). The Vendian terrane collision and accretion were accompanied by regional-scale metamorphism between 600 and 550 Ma and rejuvenation of basement igneous complexes (Korikovsky et al., 1985; Neimark et al., 1991). The onset of Late Vendian molasse deposition was in the Nikola age, when the broad shelf was formed where there was a deep water slope, and, particularly its second, advanced, stage this deposition occurred at the same time as the development of the Zherba Formation and its equivalents.
Teya-Chapa basins (Sovetov, 1993, 1997). Conglomerate at the base of the Vorogovka and Chingasan Groups contains Early Riphean metasedimentary, granite gneiss, and ophiolite bedrock pebbles (Khomichev et al., 2002). The Shangulezh conglomerate in the fill of the Iya-Tumanshet basin includes pebbles of Early Proterozoic gneiss, plagiogranite, and metasediments (Fig. 15). At the initial rifting stage, a deep-sea ramp with the Medvezh’ya coarse clastics of Early Proterozoic granites originated in place of the Teptorgo subplatform basin filled with quartzite and allite. The stage of thermal subsidence and ingression, prominent in many regions of Siberia (Igarka, the Yenisei Ridge, Sayan, and Baikal), does not show up in the deepwater environments of the Patom Upland. Recurrent rifting, inferred from the presence of overdeepened basins and deep-sea fans (Sovetov, 1997), was associated with mafic and felsic magmatism recorded in dolerite sills; in basaltic and keratophyre lavas and tuffs within the Igarka, Teya-Chapa, Iya-Tumanshet, and fore-Baikal failed rifts; and in the Patom passive margin. This rifting was coeval with the Sturtian glaciation, which produced continental and marine diamictite (tillite) in the Teya-Chapa basin (Nikolaev, 1930; Grigor’yev and Semikhatov, 1958; Sovetov and Komlev, 2005) and in the Patom passive margin (Sovetov et al., 1999). A similar origin was inferred (Shenfil’, 1991) for the chaotic breccias of the Shangulezh Formation in the lower part of the Iya-Tumanshet basin fill. The same stratigraphic level in the Vorogovka fill exhibits massive deepwater sandstone deposited by mass flows. Glaciation was completed with the development of terrigenous carbonate shelves, carbonate platforms, and carbonate turbidites. This stage of tide- and storm-influenced broad shelf (embankment) development was documented in the Sukhaya Formation in the Vorogovka basin and in the Ipsit Formation in the Iya-Tumanshet basin (Sovetov, 2001b; Blagovidov and Sovetov, 2003). The embankment facies in other basins may have been erased by Varangerian glaciers. The history of the Vendian peripheral foreland basin in the Siberian craton includes three main stages, from its inception on eroded marine fill of Late Riphean failed rifts and on the cratonic basement, to collisional culmination, and eventually to latest Precambrian plate dispersal and subsidence of the Cambrian saltbearing carbonate basin in the craton interior (Zharkov, 1974; Sovetov, 2002b; Sovetov and Blagovidov, 2004). Each stage is recorded in a second-order supersequence (Fig. 15). Marnya supersequence
LARGE SEDIMENTATION EVENTS AND STRATIGRAPHIC CORRELATION FRAMEWORK The Late Neoproterozoic geodynamic history of the southwestern craton margin of Siberia included several large tectonic events of basin evolution that were roughly coeval throughout the extensive territory. The most exact age estimates were obtained for the events in the Yenisei Ridge. Late Neoproterozoic rifting after 750 Ma is evidenced by clearly cut grabens with alluvial fans in the Vorogovka and
This supersequence (from ca. 600 Ma) encompasses the Marnya and lower Uda Formations in the Sayan region, the Aleshinsky and Chistyakov Formations, the Stolbovaya and Pod’yem Formations in the Yenisei Ridge, the Goloustnaya and Uluntui Formations in the Baikal region, the Dal’nyaya Taiga and Zhuya Groups in the Patom Upland, and possibly also the upper Chernaya Formation in the Igarka region. The lower portion of the Marnya supersequence is composed of glacial deposits and cap dolomite of the Ulyakha third-order sequence. The
Ediacaran Vendian
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Deep-sea Rhyolite and basalt lavas and tuffs, gabbro-dolerite dikes and sills A. Conglomerate at the base of riftogenic series B. Base of sincollision molasse Diamictite (tillite)
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Dz - Dzhemkukan Fm. Gr - Greben’ Fm. Gol - Goloustnaya Fm. Ips - Ipsit Fm. Kal - Kalancha Fm. Kul - Kulekin Fm. Kurt - Kurtun Fm. Kor - Korkuder Fm. Kh - Khaiverga Fm. Kha - Khorlukhtakh Fm. Lop - Lopatinsky Fm. Mr - Marnya Fm. Mar - Mariinka Fm. Med -Medvezh'ya Fm. Nik - Nikolsky Fm. Ost - Ostrovnoi Fm. Pur - Purpol Fm. Pd - Pod’yem Fm. Rd - Redkolesnaya Fm. Sh - Shangulezh Fm. St - Stolbovaya Fm. Suv - Suvorovsky Fm. Tir - Tirbess Fm. U-t - Ust’-Tagul Fm. Ugl - Uglovoi Fm. Ush - Ushakovka Fm. Vs - Veselaya Fm. Zher -Zherba Fm.
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Figure 15: Late Neoproterozoic deposition in the southwestern Siberian craton. Fm.—Formation.
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Ulyakha tillite correlates with the Varangerian and Marinoan tillites; postglacial quartzite and sandstone contain Vendian biota and correlate with the Redkino horizon of the East European Platform (Sovetov and Komlev, 2005). The glaciers eroded as the basement of the interior of the Siberian craton as Late Meoproterozoic marginal aulacogene sediment fill. The Marnya supersequence consists of three third-order sequences that bear signatures of eustatic sea level fluctuations traceable throughout the southwestern craton periphery. Clastics were transported centrifugally from the craton interior into the fringing seas during transgressions.
Greben’ time, and rifting in Ust’-Tagul time. These stages of the Vendian history of Siberia correlate with the stages of the East European Platform detected by Aks’enov (1985); the Marnya supersequence correlates with the Varanger and Redkino horizons, the Greben’ supersequence with the Kotlino horizon, and the Ust’-Tagul supersequence with the Rovno horizon. The onset of each geodynamic stage was apparently associated with plate reorganizations. The similarity of the Vendian deposition history suggests that Siberia and Baltica must have belonged to the same accretionary collage of continents within the same or neighboring climate zones.
Greben’ supersequence
DISCUSSION
This supersequence (from ca. 550 Ma) includes syncollisional continental alluvial molasse of the lower Izluchina, Tayozhny, Greben’, Veselaya, Aisa (Ikei), Khuzhir, and Bol’shoi Lug Formations, as well as fluvial deltaic and shelf deposits of the Kachergat and Zherba Formations. The lower boundary of the Greben’ supersequence is marked by the inception of alluvial plains and sharp cratonward progradation of fluvial systems. The turning point in the history of the Vendian foreland basin is especially prominent in the lower Nemchanka, Muksut, Khuzhir, and lower Kachergat fine and coarse sandstone units. The Greben’ supersequence in scrutinized sections displays four third-order sequences that record orogenic pulses and valley incision. The fill of foredeeps contains widespread redbeds of fluvial and alluvial deltaic systems and black shale prodeltas in the craton interior.
Correlation of Late Precambrian strata in the southwestern Siberian craton to the international scale has been a point of controversy (Khomentovsky et al., 1972, 1998; Sovetov, 1977, 2002c; Sovetov et al., 2003; Khomentovsky, 2002). Khomentovsky (2002) placed the terrigenous fill of foredeeps at the Upper Riphean (Baikalian or Cryogenian time) on lithostratigraphic grounds and explained the conflicting finds of latest Neoproterozoic (Vendian or Ediacaran) biota by its earlier inception in the region (Khomentovsky et al., 1998). An alternative viewpoint was originally based on correlation of sandstone petrographic suites in the Chapa, Taseeva, Oselok, and Moty Groups with sandstones from the lower platform cover stripped by exploratory petroleum wells in the craton interior (Sovetov, 1977). There the thick Late Precambrian terrigenous fill of foredeeps was found to thin and pinch out toward uplift slopes and thus to correlate with the Vendian Yudoma Group on the Aldan Shield and in the Yudoma-Maya basin (Sovetov, 1977). This subsidence and deposition pattern in the southwestern Siberian craton was accounted for by compression and lithospheric flexure that produced a system of foredeeps and forebulges (Sovetov, 2002b). Deposition in the Vendian foreland basin of the Siberian craton was discussed in detail by Sovetov and Blagovidov (2004). The latest Neoproterozoic (Vendian) age of the foreland basin, and hence its proximity in time to the Baikalian and Cadomian orogenies, was supported by the presence of tillite corresponding to the Early Varangian (Marinoan) glacial epoch at the base of the Oselok Group in the peri-Sayan foredeep (Sovetov, 2002b,c; Sovetov and Komlev, 2005). Glaciers in the Siberian craton and in the eastern and western Gondwanian cratons and microcontinents appeared at 600–590 Ma, possibly because of the formation of a large continental mass with a high-latitude position in the Southern Hemisphere (Chumakov and Sergeev, 2004; Veevers, 2004), changes in paleogeography, and inhibited oceanic water exchange. Tillites and faceted-boulder breccias found in the basal units of several formations in geographically dispersed regions (the Ulyakha member, Marnya Formation, Oselok Group, Sayan region; the Buguldeika member, Goloustnaya Formation, Baikal Group, Baikal region; and the Glubokaya member, Pod’yem Formation, Chapa Group, northeastern Yenisei Ridge)
Ust’-Tagul supersequence The lower section of this supersequence (from ca. 543 Ma) is composed of alluvial and shoreface terrigenous deposits of the Graviika, Lower Lebyazh’ya, Uglovoi, Redkolesnaya, Ust’Tagul, Shamanka (Nurtei), Ushakovka, and Kurtun Formations of the Yenisei Ridge and the Igarka, Sayan, and Baikal regions. These units were jointly interpreted as the regional YeniseiSayan Unit (Sovetov, 1977). The upper half of the supersequence includes terrigenous carbonate rocks deposited prior to the formation of the Usol’ye evaporite basin, marked by abundant salt deposits (Zharkov and Sovetov, 1969; Zharkov, 1974). The upper boundary of the supersequence coincides with the lower boundary of the Tommotian Stage of the Early Cambrian (Khomentovsky and Karlova, 2005) at ca. 535 Ma (Semikhatov, 2000). The composition of clastics in the marginal basement blocks allows inferences on the depth of erosion and on large-scale doming at the onset of the following rifting stage in the orogenic zone. The transport of voluminous mature material onto the craton was coeval to a regionally extensive transgression that produced a broad strand plain, carbonate platforms, and an evaporite basin in the Early Cambrian. The three basin evolution stages revealed in the regional supersequences reflect three main stages in the history of the Siberian craton, namely accretion in Marnya time, collision in
Sedimentary basins in the southwestern Siberian craton are identified as continental glacial deposits on the basis of sedimentology and morphology (Sovetov, 2002c; Sovetov and Komlev, 2005). The lithotype of the Ulyakha member occurs near the Ulyakha inlet into the Uda. The stacking patterns of tillite, meltwater deposits, and glaciofluvial fans reveal four stages of glacier advance. The glacier dispersed from the craton periphery and left numerous signatures of its northeastward and southwestward motion as scouring of different scales, from glacial striation to large grooves and valleys. The directions of glacier advance in the Baikal and Sayan regions are identical to those inferred from measurements in tillites of the Dzhemkukan and Nichatka Formations in the northern Patom Upland and on the western periphery of the Aldan Shield (Lungersgauzen, 1963; Chumakov, 1993). Thus, the southwestern Siberian craton must have been covered with a single ice sheet, and the glaciation may have been associated with the onset of accretion to Gondwanaland. The Ozerki member, with its cap dolomite lying over tillite, includes a set of quartzose sandstones bearing abundant metazoan trace fossils, the evidence of the earliest biota invasion of the basin during deglaciation. The remnants and traces of soft-bodied metazoan animals in the Ozerki cap dolomite and problematic fossils in the overlying terrigenous tempestite of the Bol’shaya Aisa member furnish critical evidence of the Vendian age of the tillite (Sovetov and Komlev, 2005). The Vendian age of the glaciation is supported by the fact that the typical tillite–cap dolomite association makes up a continuous regionally extensive glacial horizon with its stratigraphic position in Late Precambrian sections corresponding to the Early Varangerian stage of the Northern Hemisphere (Chumakov, 1978; Hambrey, 1983; Aitken, 1991; Narbonne et al., 1994; Arnaud and Eyles, 2002). The chemostratigraphic correlation of the Ulyakha tillite (Oselok Group, Sayan region) with δ13C variations in the Marnya and Uda carbonates confirms the Early Varangerian age of the Siberian glaciation and the Early–Late Vendian age of the metazoan biota in the Marnya Formation (Sovetov and Komlev, 2005). The δ13C curve exhibits the typical positive and negative excursions (Sovetov et al., 2003) distinguished in the model curve for Neoproterozoic III (Vendian) (Kaufman et al., 1993) and Varangerian glacial deposits (Saylor et al., 1998). The Varanger glacial epoch, including the Early and Late Varanginian stages, was accompanied by eustatic sea level fluctuations recorded in three regional sequences (Sovetov and Komlev, 2005). High-stand deposits of the Lower Uda sequence are bounded from above by the regional stratigraphic surface that traces the onset of syncollisional deposition of continental molasse (Fig. 3). This surface is marked by sandstone deposited by a deep perennial braided system with a downstream accretion complex (Foreset Macroform architectural element by Miall [1985]) of the Muksut member in the peri-Sayan foredeep and its equivalents in other foredeeps. Equivalents of the Muksut member in the craton interior were stripped by numerous petroleum wells and identified as the Bokhan member, with its lower boundary marked by the regional geophysical tiepoint M1 (Tyschenko, 1980).
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The discussion of the age of Late Precambrian molasse is intimately related to the problem of the Neoproterozoic and Vendian-Cambrian boundary. The coarse clastics of the VendianCambrian boundary are traceable almost continuously from the Igarka region as far as the Aldan Shield and make up a complex of sedimentologically similar sandstone sequences with conglomerate at the base and gradual lateral and upsection transition into terrigenous carbonate and carbonate assemblages. The identifying lithological and petrographical features of these rocks are used for tracing the Lower Cambrian (Zharkov and Sovetov, 1969) and Vendian (Khomentovsky et al., 1972; Shenfil’, 1991) or uppermost Upper Vendian (Sovetov, 2002a,b) strata limits, depending on the age assignment. Of widespread occurrence over a large area, these rocks were attributed to the basal subunit of the regional Irkut (Zharkov and Sovetov, 1969) or Yenisei-Sayan (Sovetov, 1977) horizons, according to petrographic evidence. The quartzose and arkose conglomerate and sandstone of the Graviika, Uglovoi, Redkolesnaya, Ust’-Tagul, Shamanka, Ushakovka, and upper Zherba Formations were interpreted as the top unit of Vendian collisional molasse (Sovetov, 2002a). More thorough investigation of independent lines of evidence drove us to the conclusion that this is a super-regional boundary between the tectonic stages of compression and extension on the craton (Sovetov, 2004). The stratigraphic level of the conglomerate-sandstone unit is constrained by its position immediately below the NemakitDaldyn small-shell fauna and close to the negative δ13C anomaly (Kaufman et al., 1993; Khomentovsky et al., 1998, 2004). More constraints come especially from glacial deposits at the bases of the Oselok, Chapa, and Baikal Groups and from the δ13C behavior in the Oselok carbonate correlated with that of tillite of the Blaubeker Formation from Southern Namibia (Saylor et al., 1998) and tillite at the base of the Vendian section (Sovetov et al., 2003; Sovetov and Komlev, 2005). Integrate data indicate that the conglomerate-sandstone unit of the Yenisei-Sayan horizon (Sovetov, 1977) corresponds to the Late Vendian Nemakit-Daldyn horizon of the Russian scale or to the earliest (pre-Tommotian) Lower Cambrian stage of the international scale (Gradstein et al., 2004). The deposition environments of the conglomerate-sandstone assemblage were studied by lithofacies analysis in the South Yenisei and Peri-Sayan foredeeps (Fig. 8) and on the Irkut uplift (Figs. 10 and 11). The stratotypes of the Ust’-Tagul and Shamanka Formations display the same stacking pattern as the sedimentary systems: (1) a gravelly and sandy braided-river and alternate-bar system (Figs. 8 and 10); (2) a coastal system with subaqueous sandwaves, dunes, and sheet sand with wave and current ripple structures; and (3) a coastal system with barrier islands and lagoons. The intimate interplay of continental and marine facies shows that transgression began simultaneously with orogeny, and the sea recycled an enormous mass of sand and gravel to build a broad terrigenous shelf with a minor amount of silt material. The cycles of sea progradation and back-stepping on the background of continuous transgression produced a cuspate
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shoreline dominated by either fluvial or shoreface and shelf facies. The connection with open ocean in the northern and northeastern Siberian craton is clearly detectable from terrigenous- to carbonate-facies change in the respective direction as well as from tide marks on sandwave slopes and deposits of a tidal plain that developed on a broad alluvial plain and was repeatedly drained and salinated. Lagoonal environments provided favorable conditions for metazoan detritovoral animals whose traces are found in the upper Ust’-Tagul and Shamanka sections. Examination of current directions from the dip of foresets in channel bars and sandwaves and from the orientations of flat pebbles and cobbles revealed a number of regularities. All rivers had their source in the west or south of the craton and flowed generally eastward in the northern Yenisei Ridge and in the Irkut part of the Sayan region, NNE-ward in the Biryusa and Uda-Iya parts of the Sayan region, and northwestward or northeastward in the western Baikal region. Several independent fluvial systems formed an alluvial plain and a strand plain oriented in the west-east direction and parallel to the craton edge. The alluvial plain was fringed by an epicontinental sea in the north and northeast that opened into the ocean. Fluvial sediment transport was from the orogens void of or depleted in clayey rocks, i.e., from supracrustal complexes of the uplifted craton basement blocks. The earlier inference of abrupt change in provenance of clastics at the boundary between the Aisa and Yenisei-Sayan horizons (Sovetov, 1977, 2002b) is validated by the lithology of pebbles from the Ust’-Tagul and Ushakovka conglomerates, which includes gneisses, granitoids, scarcer gabbroics, and alkaline volcanics, as well as abundant orthoquartzite, all more mature and more strongly metamorphosed than in the underlying Late
Riphean rift basins. The sediments may come from marginal basement blocks and the remnant Early Proterozoic metasedimentary fill of grabens. That the transit of rivers caused little erosion of Late Riphean and Vendian rocks is implicit evidence of the absence of folding and thrusting at that time. The quartz and feldspar-quartz (arkose) lithology and poor roundness of sandstones correspond to the “sources on stable cratons and uplifted basement” in Dickinson’s classification (1988). The orogens, which were reliably inferred to have been the source area of clastic flows, apparently were blocks of Early Proterozoic granitic basement on the western and southern craton peripheries. The rapid and roughly simultaneous uplift along more than 3000 km of the craton margin can be attributed to passive extension and doming and to local mantle flow into zones of weakness (Sovetov, 2004). The change of depositional systems in the Late Vendian– Early Cambrian conglomerate-sandstone assemblage from continental to tidal and shelf environments, from gravel (pebble) to sand fluvial and to sand shoreface facies, must have been associated with rifting. Doming and thermal subsidence were related to rapid planation of the orogens. Simultaneous rapid transgression at the boundary of the terrigenous carbonate complex of barriers and lagoons (Katanga transgression, geophysical tie M2) (Tyschenko, 1980) may have been caused by spreading and ocean opening. The clastic basin that formed at the stage of uplift framed the outer (hanging) walls of the uplifts, i.e., was a hangingwall basin. The clastic basin (alluvial fans) of the main rifting stage was destroyed by multistage Paleozoic deformation of the craton margin. The hangingwall clastic basin preceded a dramatic change in the tectonic framework and subsidence regime of the Siberian craton in the Cambrian (Fig. 16). The deposition centers moved from foredeeps to the craton interior (Zharkov, 1974). The
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Figure 16: Palinspastic reconstruction for the southwestern Siberian craton during the latest Vendian–Early Cambrian transtension and opening of the Ural Mongolian Ocean.
Sedimentary basins in the southwestern Siberian craton inception of marginal carbonate platforms and the center of noncompensated deposition in the Cambrian were responsible for the formation of an evaporation salt basin in the craton interior. The hangingwall clastic basin developed in the beginning of the tectonic stage associated with dispersal of continents, opening of transient oceanic basins, formation of passive margins, changes in oceanic circulation, and global phosphate genesis. This scenario is proved valid by the discovery of Late Vendian–Early Cambrian rift-related volcanism in the northern (Bowring et al., 1993; Pelechaty, 1996) and southwestern (Vladimirov et al., 2003) Siberian craton and outside it (Vladimirov et al., 1999). Age estimates of Neoproterozoic tectonothermal events in the Altai-Sayan and Baikal-Patom folded areas that were formerly occupied by the Paleoasian Ocean surrounding the Siberian craton are summarized by Fedotova and Khain (2002). These authors distinguish three geodynamic stages of the ocean correlated with tectonic events and basin evolution in the southwestern Siberian craton. The time interval between the onset of rifting on the Archean–Early Proterozoic craton basement and the initiation of island arc terranes and formation of subductionrelated plagiogranites is between 1050 and 1010 Ma and corresponds to the opening of the Paleoasian Ocean. Or rather it was the Proto-Paleoasian Ocean, for its initial stage was separated from the ensuing spreading (ca. 750–700 Ma) by a collisionaccretionary stage between ca. 850 and 750 Ma that produced syncollisional granites and tonalites in the Taimyr (Vernikovsky, 1996) and Baikal-Muya (Rytsk et al., 2001) fold belts and in the Tuva-Mongolia microcontinent (Kuzmichev et al., 2001). The existence of the Proto-Paleoasian Ocean is also recorded in passive margin sedimentary basins in the Yenisei Ridge (Khabarov, 1994). Accretion of terranes to the Siberian craton is proved by Late Neoproterozoic rifting (Sovetov, 1993, 1997) that acted on the accretionary complex with 750–720 Ma postcollisional granite (Vernikovsky et al., 2002). The rift basins mark the onset of the following tectonic stage, related to the breakup of Rodinia, dispersal of continents, and opening of the Paleoasian Ocean (Fig. 17). The rifts evolved into an embankment (broad shelf) in a passive regime (Sovetov, 2001b) from ca. 720 Ma to the onset of Vendian molasse deposition ca. 600 Ma. The geodynamic regime in the ocean changed in the latest Cryogenian–earliest Ediacaran (Vendian) at 640–570 Ma (Fedotova and Khain, 2002), apparently due to the effect of B-type subduction and back-arc spreading accompanied by the generation of subduction-related gabbro and granulite in the lower crust and rifting on deformed and metamorphosed Neoproterozoic complexes (Makrygina et al., 1993; Kuzmichev and Buyakaite, 1994; Vernikovsky et al., 1999; Amelin et al., 2000; Rytsk et al., 2004). At that stage, the Late Neoproterozoic failed rifts on the southwestern craton periphery experienced weak deformation and erosion. A switch to A-type subduction in the earliest Vendian (ca. 600 Ma) possibly led to accretion of island arc terranes and microcontinents to Siberia. Those events are documented in metamorphism within the Neoproterozoic terranes along the southwestern craton border and were constrained by the Ar-Ar biotite ages of 593 ± 9.7 and
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591.4 ± 3 Ma for orthogneiss and pegmatite from the Kan block and 563.3 ± 2.2 Ma for amphibolite (Nozhkin et al., 2005) and the Ar-Ar biotite ages of 554.0 ± 7 and 551.3 ± 2.9 Ma for plagiogneiss from the Arzybei, Shumikha, and Kirel terranes; 576 ± 5.8 and 546 ± 5.8 Ma Ar-Ar hornblende ages were obtained for metamorphism in the Biryusa terrane in the craton basement (Nozhkin et al., 2001). Trondjemites that intrude the Kan terrane were dated at 555 ± 5 Ma (U-Pb on zircon) (Nozhkin et al., 2001). Collision in the Patom Upland and the northern Baikal region is clearly recorded in the regional-scale metamorphism of Early and Late Neoproterozoic volcano-sedimentary and granitoid complexes. Rejuvenation of Early Neoproterozoic granites between 545 and 560 Ma (Korikovsky et al., 1985) and intrusion of the 556 ± 16 Ma Lesnaya granites (Sryvtsev et al., 1992), followed by erosion and deposition of granite pebbles in Vendian molasse (Mamakan Group), correspond to the time of Vendian tectonothermal events in the Sayan region. Chemostratigraphic correlation of the Oselok Group in the peri-Sayan foredeep with the Witvlei and Nama Groups in Southern Namibia (Sovetov and Komlev, 2005) places the accretion at 595–548 Ma and the inception of foredeeps and syncollisional molasse deposition earlier than 548 Ma (Sovetov and Blagovidov, 2004). About 570 Ma, new ophiolites of spreading centers and subduction zones formed in the Paleoasian Ocean, on the margins of the Central Mongolian and Khamar-Daban microcontinents (Fedotova and Khain, 2002). A number of dates between 544 and 522 Ma from bimodal igneous series, sills, and veins of tonalite and subalkalic gabbro record pre-Cambrian tectonic events in the ocean north and south of Siberia (Fedotova and Khain, 2002) associated with changes in plate kinematics, extension, and rifting (Pelechaty, 1996; Sovetov, 2004). Pre-Nemakit-Daldyn rifting in the northern Siberian craton is marked by swells, paleokarst features, felsic diatremes, and volcanic breccias (543 Ma) and by an unconformity at the base of the Kessyusa Formation (Bowring et al., 1993; Pelechaty, 1996). In the southwestern part of the craton, this tectonic stage was accompanied by doming, formation of alluvial fans and a clastic hangingwall basin (Sovetov, 2004), and alkali-ultramafic magmatism (Vladimirov et al., 2003). The reorganization is traceable in the Altai-Sayan folded area and in Kazakhstan from numerous Cambrian spreading centers, volcanic arcs, and related sedimentary basins (see Fig. 40 in Fedotova and Khain, 2002). The distribution and composition of magmatism may have inspired the hypothesis of the Khanty-Mansi Ocean’s opening (Şengör and Natal’in, 1996). The existence of an Early Cambrian large deep ocean that separated the Siberian craton from other continents has been proved valid by many data, including paleontological constraints (Kungurtsev et al., 2001). Spreading and volcanism were accompanied by simultaneous deposition of Early–Middle Cambrian evaporites and peripheral carbonate platforms on the craton. The evident inward migration of deposition centers from the craton periphery (Zharkov, 1974) may have been related to thermal subsidence after the Early Nemakit-Daldyn extension revealed in the northern Siberian craton (Pelechaty, 1996). The
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Figure 17: (A) Evolution of sedimentary basins in the Siberian craton, not to scale. (B) Extension of the Siberian craton in latest Vendian–Cambrian based on Wernicke’s model (1985).
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Hangingwall basin
Riphean folded marginal basins
Southwestern rifted margin
SW
Basement
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Late Riphean rift basin (failed rift) Basement
Early-Middle Cambrian deepwater basin Siberian with condensed black shales Craton Latest Vendian - Early Cambrian Late Vendian-Early-Middle Cambrian NE hangingwall basin evaporite basin Bioherm and reef Vendian foreland basin zone Delta Bar s.l. SW S S S S Vendian passive 100° margin basin
Latest Vendian-Early Cambrian dome uplift
p ga al Ba ik e La k
A
gap
gap
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Sedimentary basins in the southwestern Siberian craton formation of Cambrian sedimentary cover on cratons and microcontinents postulated by Bond et al. (1984) was coeval with the global transgression after the inception of numerous spreading centers and ocean opening (Veevers, 2004). Several dates between 510 and 450 Ma bring together granites, gneisses, trondjemites, tonalites, diorites, granulites, and, more rarely, gabbros coeval with the Middle Cambrian–Ordovician collision with Siberia of island arc and continental terranes in the Khanty-Mansi Ocean (according to Şengör and Natal’in, 1996) or the Paleoasian Ocean (according to Fedotova and Khain, 2002). The collision stage is marked in the southwestern Siberian craton by an unconformity preceding the deposition of Middle–Late Cambrian Upper Lena red molasse and the Lower Ordovician Bratsk Formation of clastics. Late Neoproterozoic tectonothermal events in the southwestern Siberian craton and the surrounding orogenic areas correlate with the evolution stages of Gondwanaland constrained by the zircon ages of 580–565 Ma obtained for the Beardmore orogen, ca. 500 Ma for the Ross-Delamerian orogens in Antarctica, and 570–530 Ma for the Avalonia and Cadomia terranes (Veevers, 2004). The earliest time span in the accretion of western and eastern Gondwanaland terranes corresponds to a brief interval of Pannotia merger (Dalziel, 1997; Veevers, 2004). The dates for the Avalonia-Cadomia orogen agree with Late Vendian collision-accretionary tectonothermal events on the western and southwestern Siberian margins. Geodynamic and climate reconstructions for the Vendian (Schettino and Scotese, 1998; Chumakov and Sergeev, 2004; Veevers, 2004) suggest that Siberia was located in the Southern Hemisphere close to Baltica and Laurentia. The collision of Baltica with Siberia, or with joined Siberia and Laurentia according to Pelechaty (1996) and Şengör and Natal’in (1996), produced the Timan and Baikal orogens. The southern and southwestern active margins of Siberia were turned toward the Avalonia-Cadomia belt and Gondwanaland, and the northern and northeastern periphery was a passive margin. The marginal position of Siberia in the tectonic collage (of Pannotia) allowed it to detach rapidly and set off in free drift during the Early Cambrian (Nemakit-Daldyn) extension associated with global plate reorganization. SUMMARY AND CONCLUSIONS The Late Neoproterozoic deposition history on the southwestern Siberian craton records four main stages of geodynamic evolution: 1. Stage I—Extension and rifting, which are especially clearly expressed in the Vorogovka and Teya-Chapa basins (failed rifts) in the Yenisei Ridge (Sovetov, 1997, 2001b) and in the Iya-Tumanshet (Karagassy) basin in the Sayan region. The relatively narrow basins, oriented at low angles to the modern craton edge, initiated on the Early Proterozoic basement and the orogenic complex of the Early Neoproterozoic passive margin. At the initial rifting stage, grabens formed and as a rule filled with
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alluvial deltaic facies along depressions. At the following stage of recurrent rifting, the rift basins were overdeepened and filled with quartz and arkose clastics transported by glaciers from the craton interior. Possible correlation with the Sturtian glacial epoch explains the relationship of rifting and climate in the context of global plate reorganization. The development of tide- and tide wave– dominated carbonate platforms is important evidence of high sea stand and spreading in the opening Paleoasian Ocean. 2. Stage II—Accretion and beginning of Paleopangea (Pannotia) assembly at 600–550 Ma, reflected in large terrigenous and carbonate shelves of remnant marginal basins of the closing Late Riphean ocean. Cratonal margins subsided under loads from overriding orogens, and the craton interior experienced the compensating uplift. The onset of this stage was coeval with the global Varanger glacial epoch and an abrupt eustatic sea level fall. The shelves of passive margins were subject to drying and exaration by glaciers in the craton periphery and interior in the earliest terminal Neoproterozoic (Early Vendian). The best representative continental tillites were found at the base of the Oselok and Baikal Groups, and their marine counterparts occur at the base of the Dal’nyaya Taiga Group in the Patom passive margin. Glaciation is marked by the till, glaciofluvial outwash plains, aeolian fields, and cap dolomite of the deglaciation stage. Representative Vendian shelf facies with subsidiary fluvial deltaic systems in the Pod’yem, Stolbovoye, Marnya, Uda, Goloustnaya, Uluntui, Valyukhta, Nikola, and Chencha Formations are spread far outside the rift basins and override deformed Meso- and Neoproterozoic deposits and the basement. The thickest shelf deposits fill foredeeps such as the East Yenisei, South Yenisei, and fore-Baikal foredeeps, which were only partly inherited from the failed rifts of the previous stage but were mostly newly initiated on a heterogeneous basement and prograded cratonward. The East Yenisei, South Yenisei, and peri-Baikal foredeeps are prominent examples of such structures. The cratonic provenance of clastics is proved by measurements of current directions in fluvial systems. This stage was the time when the Early Vendian cover began to form on the Siberian Platform. 3. Stage III—Collision (from ca. 550 Ma) and continental molasse deposition (of the Tayozhnaya, Greben’, Veselaya, Ikei, and Kachergat Formations) synchronous with tidal deposition of the Zherba Formation. Molasse filled the foredeeps and made up a belt of terrigenous deposition over 3000 km thick that was fed from the degrading surrounding orogens. The flow of clastics was so voluminous that it spilled out of the foredeeps and built the Bokhan sandstone horizon of the platform cover. 4. Stage IV—Late Vendian–Early Cambrian rifting (543– 510 Ma) associated with collapse of orogens, transtension,
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and subsidence of the craton interior with the formation of a sag basin with an evaporate zone that gradually migrated cratonward to the north and northeast. A hangingwall clastic basin formed at this stage received sediments from eroded uplifts where Early Proterozoic crystalline basement crops out. The changes in sediment sources and craton subsidence are marked by erosional and structural unconformities below the Ust’-Tagul supersequence. Thus the existence of the Paleoasian Ocean as a geodynamic system spanned an interval between ca. 750 and 570 Ma, and its reduced fragments may have persisted as late as 543 Ma. The breakup of the Vendian accretionary framework of Paleopangea (Proto-Gondwana or Pannotia) caused dramatic change to the paleogeography. The Siberian craton was surrounded by new latest Vendian–Early Cambrian oceans, including the Khanty-Mansi (Şengör and Natal’in, 1996) or Ural-Mongolian (Zonenshain et al., 1990) ocean in the southwest. Those oceans existed for a short time and began to close again in the late Middle Cambrian. ACKNOWLEDGMENTS The study was supported by grant 04-05-65299 from the Russian Foundation for Basic Research, partly by grant UR 09.01.219 from UNIROS Basic Research, and by RAS Program 6.4.1, “Geodynamic Evolution of Lithosphere in Central-Asian FoldThrust Belt (from Ocean to Continent).” REFERENCES CITED Aitken, J.D., 1991, Two Late Proterozoic glaciations, Mackenzie Mountains, north-western Canada: Geology, v. 19, p. 445–448. Aks’enov, E.M., 1985, Vendian of East-European Platform, in Vendian System, v. 2, Historic-geological and palaeontological basis: Stratigraphy and geological processes: Moscow: Nauka, p. 3–34 (in Russian). Amelin, Yu.V., Rytsk, E.Yu., Krymsky, R.Sh., Neimark, L.A., and Skublov, S.G., 2000, The Vendian age of enderbite from granulites in the BaikalMuya ophiolite belt (Northern Baikal region): U-Pb and Sm-Nd isotope evidence: Doklady RAN, v. 371, no. 5, p. 652–654. Arnaud, E., and Eyles, C.H., 2002, Glacial influence on Neoproterozoic sedimentation: The Smalfjord Formation, northern Norway: Sedimentology, v. 19, 765–788. Blagovidov, V.V., and Sovetov, Yu.K., 2003, The sedimentary structure of the Upper Precambrian Sukhaya Formation, Yenisei Ridge: Genetic analysis of Phanerozoic and Precambrian sedimentary complexes, in Yapaskurt, O.V., ed., Proceedings, 3rd Russian Lithological Workshop: Moscow: Moscow University Press, p. 205–208 (in Russian). Bond, G.C., Nickeson, P.A., and Kominz, M.A., 1984, Breakup of a supercontinent between 625 Ma and 555 Ma: New evidence and implications for continental histories: Earth and Planetary Science Letters, v. 70, no. 2, p. 325–345, doi: 10.1016/0012-821X(84)90017-7. Bowring, S.A., Grotzinger, J.P., Isachson, C.E., Knoll, A.H., Pelechaty, S.M., and Kolosov, P., 1993, Calibrating rates of Early Cambrian evolution: Science, v. 261, p. 1293–1298, doi: 10.1126/science.11539488. Bragin, S.S., 1985, The Late Precambrian Oselok Group, Sayan region: Division and correlation, in Khomentovsky, V.V., ed., Late Precambrian and Early Paleozoic stratigraphy of Siberia: Vendian and Riphean: Novosibirsk: Institute of Geology and Geophysics Press, p. 44–57 (in Russian). Bragin, S.S., 1986, Stratigraphy of the Late Riphean Karagassy Group, Sayan region: Selected problems, in Khomentovsky, V.V., ed., Late Precambrian and Early Paleozoic stratigraphy of Siberia: Stratigraphy and paleontology: Novosibirsk: Institute of Geology and Geophysics Press, p. 32–39 (in Russian).
Brasier, M.D., 1992, Global ocean-atmosphere change across the Precambrian– Cambrian transition: Geological Magazine, v. 2, p. 161–168. Butakov, E.P., Karpinsky, R.B., and Ustalov, V.V., 1975, Vendian in the TeyaChapa district: Russian Geology and Geophysics, v. 8, p. 67–77. Chumakov, N.M., 1978, Precambrian tillites and tilloids: Problems of Precambrian glacials: Moscow: Nauka, 201 p. (in Russian). Chumakov, N.M., 1993, Mid-Siberian Riphean glacial horizon: Stratigrafiya, Geologicheskaya Korrelyatsiya, v. 1, no. 1, p. 21–34. Chumakov, N.M., and Sergeev, V.N., 2004, Late Precambrian climatic zonation: Climate and biotic events, in M.A. Semikhatov and N.M. Chumakov, eds., Climate during major biotic events: Moscow: Nauka, p. 271–289 (in Russian). Dalziel, I.W.D., 1997, Neoproterozoic–Paleozoic geography and tectonics: Review, hypothesis, environmental speculation: Geological Society of America Bulletin, v. 109, p. 16–42, doi: 10.1130/0016-7606(1997)109 <0016:ONPGAT>2.3.CO;2. Dickinson, W.R., 1988, Provenance and sediment dispersal in relation to paleotectonics and paleogeography of sedimentary basins, in Kleinspehn, K.L., and Paola, C., eds., New perspectives in basin analysis (frontiers in sedimentary geology): New York: Springer, p. 3–25. Dol’nik, T.A., 2000, Stromatolity i mikrofitolity in stratigrafii pifeya I venda skladchatogo obramleniya yuga Sibirskoi Platformy (Stromatolites and microphitolites for stratigraphy of Riphean and Vendian in folded frame in the south of the Siberian Platform): Novosibirsk: GEO, 315 p. (in Russian). Fedotova, A.A., and Khain, E.V., 2002, Tectonics of the southeastern part of the Eastern Sayan and its tectonic position in the Uralian-Mongolian belt: Moscow, Scientific World, 175 p. Gladkochub, D.P., Sklyarov, E.V., Donskaya, T.V., Mazukabzov, A.M., and Stanevich, A.M., 2003, Neoproterozoic endogenic processes in the southern Siberian craton and their relation to the dispersal of Rodinia, in Sklyarov, E.V., ed., Geodynamic evolution of the lithosphere of the Central Asian foldbelt (from ocean to continent): Proceedings, Workshop on Basic Research Program, v. 1: Irkutsk: Institute of Geography Press, p. 56–59 (in Russian). Gradstein, F.M., Ogg, J.G., Smith, A.G., Bleeker, W., and Lourens, L.J., 2004, A new geological time scale, with special reference to Precambrian and Neogene: Episodes, v. 27, no. 2, p. 83–100. Grigor’ev, V.N., and Semikhatov, M.A., 1958, To question about age and genesis of so named “tillites” in northern part of Yenisey Ridge: Proceeding of USSR AS. Geol. Series, no. 11, p. 44–58. Hambrey, M.J., 1983, Correlation of Late Proterozoic tillites in the North Atlantic region and Europe: Geological Magazine, v. 120, p. 209–320. Ivanov, A.I., Lifshits, V.I., Perevalov, O.V., Strakhov, T.M., Yablonovsky, B.V., Graizer, M.I., Il’inskaya, H.G., and Golovenok, V.K., 1995, The Precambrian of the Patom Upland: Moscow, Nedra, 352 p. (in Russian). Kaufman, A.J., Jacobsen, S.B., and Knoll, F.R., 1993, The Vendian record of Sr and C isotopic variations for tectonics and paleoclimate: Earth and Planetary Science Letters, v. 120, p. 409–430, doi: 10.1016/0012-821X(93)90254-7. Khabarov, E.M., 1994, Formations and evolution of Riphean sedimentation in the eastern Yenisei Ridge: Russian Geology and Geophysics, v. 35, no. 10, p. 34–42. Khomentovsky, V.V., 2002, Baikalian of Siberia (850–650 Ma): Russian Geology and Geophysics, v. 43, no. 4, p. 299–318. Khomentovsky, V.V., and Karlova, G.A., 2005, The base of Tommotian Age— Lower boundary of Cambrian in the Siberia: Stratigrafia. Geologicheskaya Korrelyatsiya, v. 13, no. 1, p. 26–40. Khomentovsky, V.V., Shenfil’, V.Yu., Yakshin, M.S., and Butakov, E.P., 1972, Late Precambrian and Early Cambrian reference sections in the Siberian Platform: Moscow, Nauka, 356 p. (in Russian). Khomentovsky, V.V., Faizulin, M.Sh., and Karlova, G.A., 1998, The NemakitDaldyn stage of the Vendian in the southwestern Siberian Platform: Doklady RAN, v. 362, no. 6, p. 813–815. Khomentovsky, V.V., Postnikov, A.A., Karlova, G.A., Kochnev, B.B., Yakshin, M.S., and Ponomarchuk, V.A., 2004, The Vendian of the Baikal-Patom Upland, Siberia: Geologiya i Geofizika (Russian Geology and Geophysics), v. 45, no. 4, p. 465–484 (430–448). Khomichev, V.L., Kachevsy, L.K., and Amagin, A. I., 2002, Catalogue of conglomerates with igneous pebbles of Sayan-Yenisey province: Novosibirsk, SSRIGGMR, 147 p. (in Russian). Korikovsky, S.P., Sumin, L.V., and Arakelyants, M.M., 1985, Pb-Pb and K-Ar dating of Precambrian granitoids in the central Northern Baikal region (Olokit zone): Doklady AN SSSR, v. 280, no. 3, p. 688–693.
Sedimentary basins in the southwestern Siberian craton Kungurtsev, L.V., Berzin, N.A., Kazansky, A.Yu., and Metelkin, D.V., 2001, Tectonic evolution of the structure of the southwestern framing of the Siberian Platform in the Vendian–Cambrian: Paleomagnetic data: Russian Geology and Geophysics, v. 42, no. 7, p. 984–993. Kuzmichev, A.B., and Buyaka’te, M.I., 1994, Rb-Sr data of volcanics age in Shishkhid ophiolite belt (Western Peri-Khubsugul): Doklady RAN, v. 335, no. 3, p. 340–344. Kuzmichev, A.B., Bibikova, E.V., and Zhuravlev, D.Z., 2001, Neoproterozoic (~ 800 Ma) orogeny in the Tuva-Mongolian massif (Siberia): Island arc–continent collision at the northeast Rodinia margin: Precambrian Research, v. 110, p. 109–126, doi: 10.1016/S0301-9268(01)00183-8. Luchinina, V.A., Korovnikov, I.V., Sipin, D.P., and Fedoseev, A.V., 1997, Biostratigraphy of the Upper Vendian–Lower Cambrian in the Sukharikha River section (Siberian Platform): Russian Geology and Geophysics, v. 38., no. 8, p. 1346–1358. Lungersgauzen, G.F., 1963, Tillites and tilloid rocks, in Keller, B.M., ed., Stratigraphy of the USSR: Upper Precambrian, IV, v. 2: Moscow, Gostekhizdat, p. 566–577 (in Russian). Makrygina, V.A., Konnikov, E.G., Neimark, L.A., Pahol’chenko, Yu.A., Posohov, V.F., Sandimirova, G.P., Tomilenko, A.A., Tsygankov, A.A., and Vrublevskaya, T.T., 1993, The age of a granulite-charnokite complex in the Nyurundukan Formation, Northern Baikal region: A geochronological paradox: Doklady RAN, v. 332, no. 4, p. 486–490. Maslov, V.K., 1983, About breaks of sedimentation and correlation of Middle– Upper Riphean deposits in western Peri-Baikal region: Russian Geology and Geophysics, v. 24, no. 7, p. 30–42. Mats, V.D., 1962, Proterozoic rocks in the western Baikal region, in Antropov, P.Ya., ed., Geology of the USSR, XVII, Irkutsk region, geological background, v. 1: Moscow, GNTIGON, p. 79–101. Mel’nikov, N.V., 1982, Correlation of undersalt petroleum-gas bearing deposits in the south Siberian Platform: Russian Geology and Geophysics, v. 23., no. 3, p. 29–40. Miall, A.D., 1985, Architectural-element analysis: A new method of facies analysis applied to fluvial deposits: Earth-Science Review, v. 22, p. 261–308. Narbonne, G.M., Kaufman, A.J., and Knoll, A.H., 1994, Integrated chemostratigraphy and biostratigraphy of the Windermere Supergroup, northwestern Canada: Implications for Neoproterozoic correlation and the early evolution of animals: Geological Society of America Bulletin, v. 106, no. 10, p. 1281–1292. Neimark, L.A., Rytsk, E.Yu., Gorokhovsky, B.M., Ovchinnikova, G.V., Kiseleva, E.I., and Konkin, V.D., 1991, Pb isotope composition and the origin of Pb-Zn mineralization in the Olokit zone, Northern Baikal region: Geologiya Rudnyh Mestorozhdenii, v. 33, no. 6, p. 34–49. Nikolaev, I.G., 1930, Glacial deposits (tillites) of Lower Cambrian age in the Yenisei Ridge: Bulletin of the Geological and Prospecting Service of the U.S.S.R., v. 49, no. 7, p. 1–15 (in Russian). Nozhkin, A.D., Turkina, O.M., Bibikova, E.V., and Ponomarchuk, V.A., 1999, Riphean granite gneiss domes of the Yenisei Range: Geological structure and U-Pb isotopic age: Russian Geology and Geophysics, v. 40, no. 9, p. 1284–1292. Nozhkin, A.D., Turkina, O.M., Bibikova, E.V., and Ponomarchuk, V.A., 2001, Structure, composition, and formation conditions of metasedimentary-volcanogenic complexes of the Kan greenstone belt (northwestern cis-Sayan region): Russian Geology and Geophysics, v. 42, no. 7, p. 999–1019. Nozhkin, A.D., Sovetov, Yu.K., Travin, A.V., and Turkina, O.M., 2005, Geological and isotope geochronology evidence for Vendian collision-accretionary events in the southwestern margin of the Siberian craton, in Karyakin, Yu.V., ed., Crustal and mantle tectonics: Tectonic control on distribution of mineral deposits: Proceedings, XXXVIII Tectonic Workshop, v. 2: Moscow, GEOS, p. 52–55 (in Russian). Pelechaty, S.M., 1996, Stratigraphic evidence for the Siberia-Laurentia connection and early Cambrian rifting: Geology, v. 24, p. 719–722, doi: 10.1130/0091-7613(1996)024<0719:SEFTSL>2.3.CO;2. Puchkov, V.N., 2001, Have the Late Vendian supercontinents ever existed? in Sklyarov, E.V., ed., Supercontinents in the Precambrian geologic history: Proceedings, Scientific Workshop: Irkutsk, Institute of the Earth’s Crust, p. 213–218 (in Russian). Rytsk, E.Yu., Amelin, Yu.V., Rizvanova, N.G., Krymsky, R.Sh., Mitrofanov, G.L., Mitrofanova, N.N., Perelyaev, V.I., and Shalaev, V.S., 2001, The age of rocks in the Baikal-Muya fold belt: Stratigrafiya, Geologicheskaya Korrelyatsiya, v. 9, no. 4, p. 3–15.
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Rytsk, E.Yu., Makeev, A.F., Glebovitsky, V.A., and Fedoseenko, A.M., 2004, The Vendian (590Þ±5 Ma) age of the Padry Group rocks in the Baikal-Muya fold belt: U-Pb zircon data: Doklady RAN, v. 397, no. 4, p. 517–519. Rzhevsky, V.F., and Chekhovich, K.M., 1978, Precambrian stratigraphy of the Igarka region, in V.V. Khomentovsky, ed., New evidence in Late Precambrian stratigraphy and paleontology of eastern and northern Siberia: Novosibirsk: Institute of Geology and Geophysics, p. 101–113 (in Russian). Saylor, B.Z., Kaufman, A.J., Grotzinger, J.P., and Urban, F., 1998, A composite reference section for terminal Proterozoic strata of Southern Namibia: Journal of Sedimentary Research, v. 68, no. 6, p. 1223–1235. Schettino, A., and Scotese, C.R., 1998, Paleo-continental Map Editor for Windows NT 4.0 version, CD-ROM: Arlington, University of Texas. Semikhatov, M.A., 1962, Riphean and Lower Cambrian stratigraphy of the Yenisei Ridge: Moscow, Nauka, 242 p. (in Russian). Semikhatov, M.A., 2000, Refining isotope dates of the lower boundary of the Upper Riphean, Vendian, Upper Vendian, and Lower Cambrian, in Zhamoida, A.I., ed., Supplements to the Stratigraphic Code: St. Petersburg: VSEGEI, p. 95–107 (in Russian). Şengör, A.M., and Natal’in, B.A., 1996, Turkic-type orogeny and its role in the making of the continental crust: Annual Review of Earth Science and Planetary Sciences, v. 24, p. 263–337, doi: 10.1146/annurev.earth.24.1.263. Shenfil’, V.Yu., 1991, Late Precambrian stratigraphy of the Siberian Platform: Novosibirsk, Nauka, 183 p. (in Russian). Shishkin, B.B., 1975, Stratigraphy of Cambrian and Pre-Cambrian bounding layers from Igarka region, in Sokolov, B.S., and Khomentovksy, V.V., eds., Analogues of Vendian complex in Siberia: Proceedings of the Conference on the Vendian Stratigraphy of the Siberian Platform: Transactions, Institute of Geology and Geophysics, v. 232: Moscow, Nauka, p. 192–196 (in Russian). Sovetov, J.K., 1977, Upper Precambrian sandstones in the southwestern Siberian Platform: Novosibirsk, Nauka, 294 p. (in Russian). Sovetov, J.K., 1993, Sedimentologic sequences in Late Riphean aulacogens at south-western margin of the Siberian Platform related to sea-floor spreading, in Dobretsov, N.L., and Berzin, N.A., eds., IGCP Project 283, Report no. 4, Fourth International Symposium on Geodynamic Evolution of Paleoasian Ocean, Abstracts: Novosibirsk: UIGGM, p. 164–167. Sovetov, J.K., 1997, Late Riphean rifting and the Baikal geodynamic cycle of the Siberian Platform, in Koroteev, V.A., ed., The Riphean of northern Eurasia: Geology and general problems of stratigraphy: Ekaterinburg, Institute of Geology and Geochemistry UB RAS, p. 223–230. Sovetov, J.K., 2001a, Geodynamic evolution of Siberian and East-European cratons, comparison by analysis of Vendian foreland basins, in Letnikov, F.A., ed., Geology, geochemistry, and geophysics at the turn of the 20thto-21st century: Proceedings, Russian Conference: Irkutsk, Institute of the Earth’s Crust, p. 120–122 (in Russian). Sovetov, J.K., 2001b, A Vendian peripheral foreland basin of the Siberian craton: Stages of Paleopangean accretion, in Supercontinents in the Precambrian geologic history: Proceedings, Scientific Workshop: Irkutsk, Institute of the Earth’s Crust, p. 254–258 (in Russian). Sovetov, J.K., 2001c, The Vorogovka Group: Evolution of a typical failed rift basin associated with the dispersal of Rodinia (Upper Riphean, Yenisei Ridge), in Supercontinents in the Precambrian geologic history: Proceedings, Scientific Workshop: Irkutsk, Institute of the Earth’s Crust, p. 258–262 (in Russian). Sovetov, J.K., 2002a, Vendian foreland basin of the Siberian cratonic margin: Paleopangean accretionary phases: Russian Journal of Earth Sciences, v. 4, no. 5, p. 363–387. Sovetov, J.K., 2002b, Vendian glaciation on the Siberian craton, in Letnikov, F.A., Geology, geochemistry, and geophysics at the turn of the 20th-to21st century: Proceedings, Russian Conference: Irkutsk, Institute of the Earth’s Crust, p. 122–124 (in Russian). Sovetov, J.K., 2004, The doming stage of rifting and breakup of Paleopangea (Pannotia) in the Late Vendian–Early Cambrian: Evidence from a peripheral foreland basin on the southwestern Siberian craton, in Leonov, Yu.G., Evolution of tectonic movements in the Earth’s history: Proceedings, Russian Conference, v. 2: Novosibirsk, GEO, PH SB RAS, p. 192–194 (in Russian). Sovetov, J.K., and Blagovidov, V.V., 2004, Reconstruction of a basin evolution: Example from a Vendian foredeep, a foreland basin in the southwestern Siberian craton, in Leonov, Yu.G., and Volozh, Yu.A., eds., Sedimentary basins: Structure, evolution, and methods of studies: Moscow: Nauchny Mir, p. 159–212 (in Russian). Sovetov, J.K., and Blagovidov, V.V., 2006, Late Riphean deposition in the Iya-Tumanshet autacogene (southwestern of the Siberian platform):
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Supersequences and correlation with the Riphean stratotype in the Bashkiriya anticlinorium, in Koroteev, V.A., ed., Litologicheskie aspekty sloistykh sred v Geologii, Proceedings of 7th Ural Lithological Meeting: Ekaterinburg, Institute of Geology and Geochemistry of the UB RAS, p. 248–250 (in Russian). Sovetov, J.K., and Komlev, D.A., 2005, Tillite at base of the Oselok Group in the Sayan region and the position of the lower boundary of the Vendian in the southwestern Siberian Platform: Statigrafiya, Geologicheskaya Korrelyatsiya, v. 13, no. 4, p. 3–34. Sovetov, J.K., Akul’shina, E.P., Ivanovskaya, A.V., and Pisareva, G.M., 1975, Stacking pattern, lithology, and deposition environments of the Yudoma complex in the southeastern Yenisei Ridge, in Akul’shina, E.P., and Ivanovskaya, A.V., eds., Lithological and geochemical studies of Paleozoic and Precambrian strata: Novosibirsk, Institute of Geology and Geophysics Press, p. 81–100 (in Russian). Sovetov, J.K., Khabarov, E.M., Blagovidov, V.V., and Saraev, S.V., 1999, The Riphean divergent continental margin of Siberia (Patom Upland): Basin evolution, in Geological history of Proterozoic pericratonal and paleooceanic structures in Northern Eurasia: Proceedings, Russian Conference: St. Petersburg, Tema, p. 159–162. Sovetov, J.K., Ponomarchuk, V.A., and Komlev, D.A., 2003, The Vendian lower boundary and stratigraphy in the southwestern Siberian Platform, from δ13 C variations in carbonates, in Kozakov, I.K., and Kotov, A.B., eds., Isotope geochronology for geodynamics and metallogeny: Proceedings, 2nd Russian Conference on Isotope Geochronology: St. Petersburg, Tsentr Informatsionnoi Kul’tury, p. 473–476 (in Russian). Sryvtsev, N.A., Khalilov, V.A., Buldygerov, V.V., and Perelyaev, V.I., 1992, Geochronology of granitoids of the Baikal-Muya belt: Russian Geology and Geophysics, v. 33, no. 9, p. 60–64. Tetyaev, M.M., 1916, To geology of the western Peri-Baikal, in Materials to historical and applied geology, Part 2: Petrograd, Geological Committee, 55 p. (in Russian). Turkina, O.M., Bibikova, E.V., and Nozhkin, A.D., 2003, Stages and environments of Early Proterozoic plutonism in the southwestern margins of the Siberian craton: Doklady RAN, v. 388, no. 6, p. 779–783. Tyschenko, L.F., 1980, The Moty Formation in the Irkutsk amphitheater: Regional correlation, in Karogodin, Yu.N., ed., Problems of lithostratigraphy: Novosibirsk, Nauka, p. 149–158 (in Russian). Veevers, J.J., 2004, Gondwanaland from 650–500 Ma assembly through 320 Ma merger in Pangea to 185–100 Ma breakup: Supercontinental
tectonics via stratigraphy and radiometric dating: Earth Science Reviews, v. 68, 132 p. Vernikovsky, V.A., 1996, Geodynamic evolution of the Taimyr folded area: Novosibirsk, UIGGM, 203 p. Vernikovsky, V.A., Vernikovskaya, A.E., Sal’nikova, E.B., Kotov, A.B., and Chernyh, A.I., 1999, New U-Pb data on the formation of the Predivinsk paleoisland-arc complex (Yenisei Range): Russian Geology and Geophysics, v. 40, no. 2, p. 256–261. Vernikovsky, V.A., Vernikovskaya, A.E., Sal’nikova, E.B., Kotov, A.B., and Kovach, V.P., 2002, Postcollisional plutonism in Zaangar’ye, Yenisei Ridge: Events between 750 and 720 Ma: Doklady RAN, v. 384, no. 2, p. 221–226. Vernikovsky, V.A., Vernikovskaya, A.E., Kotov, A.B., Salnikova, E.B., and Kovach, V.P., 2003, Neoproterozoic accretion-collisional events on the western margin of the Siberian Craton: New geological and geochronological evidence from the Yenisey Ridge: Tectonophysics, v. 375, p. 147–168, doi: 10.1016/S0040-1951(03)00337-8. Vladimirov, A.G., Ponomareva, A.P., and Kargopolov, S.A., 1999, Neoproterozoic ages of the oldest rocks of the Tomsk uplift (Gornaya Shoria), from U-Pb, Sm-Nd, Rb-Sr, and Ar-Ar data: Stratigrafiya, Geologicheskaya Korrelyatsiya, v. 7, no. 5, p. 28–42. Vladimirov, B.M., Logachev, N.A., Vainer-Krotova, G.A., Lepin, V.S., Ivanov, A.V., and Rasskazov, S.V., 2003, The Vendian/Cambrian boundary: RbSr isochron ages of the final event of calc-ultramafic magmatism in the Sayan region: Doklady RAN, v. 389, no. 6, p. 777–780. Wernicke, B., 1985, Uniform sense simple shear of the continental lithosphere: Canadian Journal of Earth Sciences v. 22, p. 108–125. Zharkov, M.A., 1974, Paleozoic salt-bearing formations of world: Moscow, Nedra, 391 p. Zharkov, M.A., and Sovetov, Yu.K., 1969, Irkut horisons, their extent and stratigraphic position, in Sokolov, B.S., ed., Lower Cambrian and Upper PreCambrian stratigraphy of Siberian Platform south: Transactions, Institute of Geology and Geophysics, v. 51: Moscow, Nauka, p. 34–53 (in Russian). Zonenshain, L.P., Kuz’min, M.I., and Natapov, L.M., 1990, Geology of the USSR: A plate tectonic synthesis, in Page, B.M., ed., American Geophysical Union (AGU) Geophysical Monograph, Geodynamics Series 21, Washington, D.C., AGU.
MANUSCRIPT ACCEPTED BY THE SOCIETY 3 OCTOBER 2006
Printed in the USA
Geological Society of America Special Paper 423 2007
Aluminum phosphate in Proterozoic metaquartzites: Implications for the Precambrian oceanic P budget and development of life Giulio Morteani* Gmain No. 1, 84424 Isen, Germany Dietrich Ackermand Mineralogisch-Petrographisches Institut, Christian-Albrechts-Universität Kiel, Olshausenstrasse 40, 24105 Kiel, Germany Jörg Trappe Geologisches Institut, Universität Bonn, Nussallee 8, 53115 Bonn, Germany ABSTRACT The present article reports the textures, whole-rock and mineral chemistry, and mineral reactions under greenschist- to amphibolite-facies conditions of Proterozoic aluminum phosphate–bearing metaquartzites from central Madagascar and from the Espinhaço fold belt in Brazil. Based on this information, the mechanism of aluminum phosphate genesis in the protolith of the metaquartzites and the behavior of phosphates (lazulite, augelite, trolleite, svanbergite, goyazite-crandallite, berlinite, and xenotime) during low-grade amphibolite-facies metamorphism is discussed. In Precambrian marine sandstones, aluminum phosphate was a potentially important phosphorus sink before the beginning of abundant organic and large-scale calcium phosphate deposition in the Neoproterozoic. A compilation of alternative calcium phosphate (apatite) deposits in the Precambrian confirms their insignificance for phosphorus fixation before the Neoproterozoic. Efficient phosphorus burial in aluminum phosphates inhibited the rapid buildup of phosphorus nutrient levels in the early Precambrian oceans. Keywords: aluminum phosphates, apatite, phosphate sediments, phosphorites, bioproduction, Brazil, Madagascar, metamorphism, ocean, phosphorus, Precambrian
INTRODUCTION Although phosphorus is not an extremely rare element in the Earth’s crust today, it is a critical element in the biosphere. Phosphorus is an essential nutrient for all forms of life and is the limiting factor for bioproduction. In modern oceans roughly 85% of phosphorus buried in the marine environment is bound to organic calcium phosphate *E-mail:
[email protected].
deposition (Filippelli, 2002). In the early Proterozoic oceans, a high phosphorus content was caused by the small amount of organic matter in the biosphere (Des Marais et al., 1992) at that time and the consequent lack of widespread organic calcium phosphate precipitation. Due to the secondary importance of organically driven calcium phosphate formation, the precipitation of aluminum phosphate minerals in sandstones played an important role from Middle Archean to Early Cretaceous times (Rasmussen, 1996). The phosphates identified by Rasmussen (1996) in the unmetamorphosed sandstones of Western Australia are xenotime,
Morteani, G., Ackermand, D., and Trappe, J., 2007, Aluminum phosphate in Proterozoic metaquartzites: Implications for the Precambrian oceanic P budget and development of life, 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. 579–592, doi: 10.1130/2007.2423(29). For permission to copy, contact
[email protected]. ©2007 Geological Society of America. All rights reserved.
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florencite, crandallite, gorceixite, and apatite. Marine sandstones are a typical and common lithology of the northern margin of the supercontinent Gondwana, and not only in the Proterozoic. Siliciclastic deposystems persisted during the early Paleozoic rifting and opening of the Rheic Ocean. Most likely, aluminum phosphate precipitation as documented in the present article for Proterozoic marine sandstones continued to be a significant phosphorus sink during the opening of the Rheic Ocean after the beginning of abundant organic deposition and the consequent large-scale calcium phosphate formation in the Neoproterozoic. Precambrian rocks are typically metamorphosed. Aluminum phosphate–bearing (APB) metaquartzites generated from Proterozoic marine phosphate-rich sandstones by metamorphism have rarely been described in the literature in spite of the fact that aluminum phosphate minerals—due to their wide stability field, which ranges from diagenesis (Spötl, 1990; Rasmussen, 1996; Rasmussen et al., 1998) to amphibolite facies (Wise and Loh, 1976; Schmid-Beurmann, 1997; Schmid-Beurmann et al., 1997; Morteani and Ackermand, 2004), and their refractory nature—record the whole geological history of the host rocks, from sedimentation to their metamorphic overprint. The rarity of APB quartzites is very likely apparent only due to and produced by the fine-grained nature of the phosphates crystallized in the intergranular spaces that are not recognizable by conventional microscopy. In contrast to the APB metaquarzites, Precambrian discrete phosphate sediments and phosphorites were systematically investigated in the past due to their potential economic significance (Cook and Shergold, 1986a,b). Although the facies and depositional environments of Precambrian phosphate sediments differ widely, the time distribution of larger occurrences is bound to the later Proterozoic or Ediacaran. This led to the conclusion that a phosphorus sink was lacking in the early Precambrian ocean and that the ocean water was consequently rich in phosphorus (Einsele, 1992). The early and efficient burial of phosphorus via aluminum phosphate precipitation slowed down the buildup of the phosphorus nutrient level in the water of the Proterozoic oceans that was needed for the development of abundant life until the beginning of the Neoproterozoic. After the Late Proterozoic breakup of the supercontinent and the rising of sea levels at the end of the Neoproterozoic glaciations, the opening Iapetus and Rheic oceans were surrounded by very broad continental shelves (Golonka et al., 1994). The northern margin of Gondwana in particular remained continuously flooded far inland until the demise of the Rheic ocean in the Carboniferous. The shallow marine areas provided the environments for the evolution of the metazoans but also for widespread siliciclastic depositional systems with potential aluminum phosphate formation. No specific information exists about the phosphate mineralogy in psammitic sediments deposited in the Rheic Ocean. The psammitic sediments with their phosphate minerals should give key information on phosphogenesis along the northern margin of Gondwana and in the Rheic Ocean from the Neoproterozoic to the Carboniferous. The present article summarizes and discusses literature data and our own results (Morteani and Ackermand, 2001,
2004, 2006) on the texture, chemical and mineralogical composition, and protolith of the Proterozoic APB metaquartzites found in the Serra do Espinhaço (Minas Gerais, Bahia, Brazil) and southern Madagascar as evidence of the Proterozoic marine phosphogenesis and of a major phosphorus sink in the Proterozoic oceans. In the Proterozoic, APB metaquartzites played a key role in phosphogenesis as compared to that of the rather rare Proterozoic apatite-bearing sediments. GEOLOGICAL SETTING AND EXISTING DATA Metaquartzites The localities of the studied APB metaquartzites and phosphorite-bearing rocks are given in Figure 1. The studied APB metaquartzites of central Madagascar belong to the Itremo Group (Lacroix, 1922–1923; Cox et al., 1998; Fernandez and Schreurs, 2003; Morteani and Ackermand, 2006). The sedimentation age of the protolith of the metaquartzites is between 2000 and 1700 Ma (Fernandez and Schreurs, 2003). The sedimentary textures and mineral parageneses of the sandstones were overprinted by the Pan-African tectonothermal event between 780 and 570 Ma (Collins et al., 2003; Morteani and Ackermand, 2006). The APB metaquartzites of the Serra do Espinhaço were deposited in the Espinhaço rift within the São Francisco craton (Bruni and Schobbenhaus Filho, 1976). The Espinhaço rift started to open in the late Paleoproterozoic at ca. 1750 Ma (Almeida Abreu, 1993). In the Mesoproterozoic, between ca. 1350 and 1000 Ma the rift closed and the protolith of the metaquartzites was overprinted, first by the Uruaçuano-Espinhaço event (Brito Neves et al., 1979; López and De Souza, 1985; Kalt, 1991; Cordani et al. 1992; Almeida Abreu, 1993) and then at ca. 485 ± 25 Ma by the Brasiliano tectonometamorphic event (Tavora et al., 1967). As documented by well-preserved primary sedimentary structures such as wave ripples, planar lamination, and crossbedding, all studied APB metaquartzites were deposited in a shallow continental shelf environment (Moine, 1967; Cox et al., 1998; Morteani and Ackermand, 2001, 2004, 2006). In zones of intense tectonic deformation, sedimentary structures can be completely obliterated. According to Bhatia (1983), sandstones show decreasing sums of (Fe2O3 + MgO) of the ratio Al2O3/SiO2 and of TiO2 contents going from oceanic to continental arcs, active margins, and passive margins (Fig. 2A and B). In the (Fe2O3 + MgO) versus TiO2 and the (Fe2O3 + MgO) versus Al2O3/SiO2 diagrams, the chemical composition of the studied APB metaquartzites plots with low (Fe2O3 + MgO), TiO2, and Al2O3/SiO2 data in the field of the passive margin, indicating a depositional environment on a passive margin. Phosphorites A total of twenty-four Precambrian economic and noneconomic phosphate deposits are listed and systematically
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© CP Figure 1. World map showing the locations of mentioned aluminum phosphate–bearing (APB) metaquartzite and phosphorite deposits. APB metaquartzites: A—Espinhaço fold belt, Brazil; B—Itremo Group, central Madagascar. Phosphorites: 1—Patos de Minas, Brazil, Neoproterozoic (? Ediacaran); 2—Upper Volta, Niger, Ediacaran; 3—Lake Khubsugul, Mongolia, Ediacaran; 4—Udaipur, India, Paleoproterozoic; 5—Singpung, North Korea, Paleoproterozoic; 6—South Siberia / Lake Baikal, Russia, Ediacaran; 7—Yangtze Platform, southwestern China, Ediacaran; 8—Adelaide geosyncline, Australia, Neoproterozoic; 9—Amadeus basin, Australia, Neoproterozoic; 10—Krol Valley, northern India, Late Proterozoic; 11—Ciudad Real, Spain, Ediacaran; 12—Lampinsaari, Finland, Paleoproterozoic; 13—Temo, Finland, Paleoproterozoic; 14—Värmland, Sweden, Paleoproterozoic; 15—Lake Vättern, Sweden, Paleoproterozoic; 16—Cailleach Head, (northern UK, Scotland), Neoproterozoic; 17—Wollaston Lake fold belt, northern Canada (Saskatchewan), Paleoproterozoic; 18—Penokean fold belt, USA (Michigan, Minnesota), Paleoproterozoic; 19—Thelon basin, northern Canada (Northwest Territories), Mesoproterozoic; 20—Athabasca basin, northern Canada (Saskatchewan), Mesoproterozoic; 21—northern Rocky Mountains, Belt Series, USA (Montana), Mesoproterozoic; 22—southern Appalachians, Ocoee Supergroup, USA (Tennessee, North Carolina), Neoproterozoic (? Ediacaran); 23—Vendougou, Senegal, Ediacaran; 24—High Atlas, Morocco, Ediacaran.
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described by an international research group in the framework of the International Geological Correlation Program (IGCP 156: Phosphorites) (Cook and Shergold, 1986a) (Fig. 1). In addition, most banded ironstone formations contain sporadic phosphate bands or pebbles. North America Phosphate pebbles were reported from the Paleoproterozoic Marquette Range Supergroup in the Penokian fold belt in Michigan, and thin phosphate beds were described from graphitic phyllites and quartzites of the lateral equivalent, the Thomson Formation in Minnesota (Fig. 1, 18). Paleoproterozoic sedimentary apatite-uranium mineralizations occur in the Wollaston Lake fold belt in Saskatchewan (Fig. 1, 17). Mesoproterozoic phosphate sediments comprise apatite grains and cements in the Wolverine Point Sandstone in the Athabasca basin (Fig. 1, 20) and further apatite and goyacite cements in the Thelon Quartzite of the Canadian Thelon basin (Fig. 1, 19). Thin layers of grainstone phosphorite occur in the Mesoproterozoic Belt Supergroup in Montana (Fig. 1, 21), and phosphate nodules occur in the Neoproterozoic, probably Ediacaran, Wilhite Formation (Walden Creek Group) of the upper Ocoee Supergroup in Tennessee (Fig. 1, 22) (Christie and Sheldon, 1986). All these deposits are minor and of noneconomic scale. South America Economic phosphorites were discovered in the Neoproterozoic on the basis of underlying glaciogenic sediments, dated possibly as part of the Ediacaran Upper Bambui Group (Fig. 1, 1). The granular phosphate beds, phosphatic shales, and phosphate cementations are associated with partly biolaminated carbonates and shales (Cathcart, 1977; Dardenne et al., 1986). Europe Precambrian phosphate deposits are only sporadically noted. On the Baltic Platform at Lampinsaari and Temo in Finland and also at Värmland in Sweden (Fig. 1, 12, 13, and 14), apatite lenses and nodules occur in Paleoproterozoic metamorphic dolomites and gneisses and are considered to originate from a sequence of acidic tuffs. Ek and Nysten (1990) describe aluminum phosphate mineralization from the same general area. At Lake Vättern in Sweden, flattened phosphate nodules and phosphate pebbles are present in the quartzites or conglomerates of the “sparagmites,” a thick deltaic or fluvial succession of Paleoproterozoic age (Fig. 1, 15) (Notholt and Brasier, 1986). In Scotland, phosphate nodules and lenses in shales within the Neoproterozoic Torridonian Group are mentioned (Fig. 1, 16). Grainstone phosphorites from a pelitic and siliciclastic sequence in Central Spain (Fig. 1, 11) with a possible Vendian age are described by Gabaldón López et al. (1989). All occurrences in Europe are minor and of noneconomic scale. Africa The Ediacaran pelitic-siliciclastic fill of the Volta basin (Fig. 1, 2) in Burkina Faso, Niger, and Benin hosts potentially
economic pelletal and ooidic phosphorites at the base (Trompette, 1989). Minor occurrences of similar age have been described from the High Atlas Mountains of Morocco and from eastern Senegal (Slansky, 1986) (Fig. 1, 23 and 24). Australia Small noneconomic deposits were described from two levels in the Neoproterozoic Adelaide geosyncline, consisting of granular and ooidic phosphorites and further nodules and lenses (Fig. 1, 8). In the Amadeus basin, glauconitic-phosphatic shales and sandstones as well as phosphatic mudstone bands occur in the Bitter Springs, Areyonga, and Pertatataka Formations, of Neoproterozoic age (Howard, 1986) (Fig. 1, 9). Asia Phosphate deposits are known from Korea, China, India, and the southern Siberian Platform. All of these are significant economic resources. The Mach’onnyang Series in North Korea comprises marbles, mica and graphite schists, amphibolites, and other paragneisses, including phosphate lenses (Fig. 1, 5). With its Paleoproterozoic age, this deposit is among the oldest phosphate sediments (Notholt et al., 1989). Significant and widespread mudstone and granular phosphorites occur in the Ediacaran Doushantuo Formation and equivalents on the Yangtze Platform of southwestern China (Li, 1986) (Fig. 1, 7). Among other minor phosphate deposits, significant phosphorites exist in Rajasthan and Madhya Pradesh in Central India and in the outer Himalayan belt (Fig. 1, 4). The Paleoproterozoic Aravalli Group (older than 1800 Ma) in Central India, with a major deposit at Jhamarkotra, consists of stromatolitic carbonates, quartzites, and phyllites, with the stromatolites widely phosphatic. In the Himalayan region, the Upper Krol Group, of Neoproterozoic to ?Cambrian age, comprises mudstone and granular phosphate layers or lenses within a black shale and dolomite suite (Fig. 1, 10). In the same general region, the Tal Formation at the Mussoorie syncline, formerly assigned a Vendian age, is now considered Cambrian and consists of chert and black shale with phosphate mudstone and phosphatic intraclasts (Banerjee, 1986; Banerjee et al., 1986). The Khubsgul basin in northern Mongolia (Fig. 1, 3) hosts a suite of intraclastic phosphorites in the lower, Ediacaran, portion of the calcareous Khubsgul Group (Ilyin et al., 1989). The deposits continue into South Siberia, where farther west, in the Vendian Belka Suite of the Altay-Sayan, brecciated stromatolitic dolomites are phosphatized (Krasilnikova and Ilyin, 1989) (Fig. 1, 6). Archean phosphorites or phosphate sediments are unknown. Calcium phosphate deposition began in the Early Proterozoic in very minor occurrences except for the North Korean deposit and the stromatolitic deposits in India. In the Neoproterozoic, specifically in the Ediacaran, the first large-scale deposits were formed as part of the Neoproterozoic–Cambrian phosphorite giant. The stratigraphic and volumetric data of most Precambrian phosphate deposits are too sparse for a reliable quantification of the phosphorus flux into the Precambrian sediments.
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TABLE 1. SELECTED MAJOR AND TRACE ELEMENT WHOLE-ROCK COMPOSITIONS OF ALUMINUM PHOSPHATE–BEARING METAQUARTZITES FROM MADAGASCAR AND BRAZIL Sample no. Madagascar Brazil 12953 10808 5868B 13567 13589 5726 5714 5715 Percent SiO2 90.31 91.30 77.69 72.43 82.50 62.28 63.80 68.30 0.08 0.07 0.17 0.14 0.06 0.07 0.09 0.25 TiO2 3.90 3.80 10.10 9.23 5.56 16.99 21.00 16.10 Al2O3 1.54 1.76 0.29 0.38 0.28 1.84 2.38 3.47 Fe2O3 MnO 0.0011 0.0008 0.0009 0.0025 0.0017 b.d.l. b.d.l. b.d.l. MgO <0.01 0.08 0.21 2.74 1.78 2.41 0.70 0.70 CaO 0.08 0.07 0.09 0.16 0.09 0.19 0.02 0.02 0.01 <0.02 0.08 0.02 0.02 0.14 0.14 0.06 Na2O K2O 0.05 0.09 0.06 0.16 0.07 1.79 5.38 3.48 P2O5 1.91 1.13 9.37 10.73 6.73 11.69 3.55 4.64 LOI 0.90 0.60 1.60 2.00 1.30 3.50 2.60 2.80 Total 98.80 98.90 100.00 98.30 98.60 101.63 99.70 99.82 Parts per mil Ba b.d.l. 26 b.d.l. 24 Rb b.d.l. 5 b.d.l. 4 S 700 400 b.d.l. b.d.l. Sr 5581 3142 1153 2567 Th b.d.l. 5 5.5 5.6 U b.d.l. b.d.l. 0.9 b.d.l. Y 95 74 144 249 Zr 221 143 389 365 B b.d.l. b.d.l. b.d.l. 13 Notes: b.d.l.—below detection limit; LOI—loss on ignition.
GEOCHEMISTRY Quartzites Whole-Rock Major and Trace Element Composition The whole-rock chemical compositions of APB metaquartzite samples selected as representative from Morteani and Ackermand (2001, 2004, 2006) are given in Table 1. The analytical methods are explained by Morteani and Ackermand (2001, 2004). In the CaO versus P2O5 plot, two point clusters can be distinguished (Fig. 3A). One cluster is characterized by P2O5 contents of 2 wt% (apatite) maximum. The other cluster is given by samples with a P2O5 content up to 12 wt% (aluminum phosphates). In the cluster marked as apatite, all samples from Madagascar are found; the cluster marked aluminum phosphates encompasses predominantly samples deposited in Brazil in the Espinhaço rift. The P2O5 to CaO ratio in the Brazilian samples is by far higher than that given by the P2O5 to CaO ratio in apatite, which is ~0.7. This indicates that in the protolith of the metaquartzites from Brazil, apatite was newer, or only very subordinately the major phosphate carrier, whereas for the second cluster the possibility of apatite as a primary sedimentary mineral cannot be excluded. For the samples from Madagascar, the CaO versus Sr plot shows a positive correlation due to the typical substitution of Ca by Sr shown by apatite (Altschuler, 1980) and aluminum phosphates such as the goyazite-crandallite solid solution series (Fig. 3B). The Brazilian samples show a large scatter without any
69 b.d.l. b.d.l. 612 b.d.l. b.d.l. b.d.l. 136 b.d.l.
461 52 b.d.l. 1646 b.d.l. b.d.l. 75 51 b.d.l.
153 339 149 88 0.04 b.d.l. 476 4889 b.d.l. b.d.l. b.d.l. b.d.l. 43 54 b.d.l. 116 477 14
correlation. The scatter is due to the presence of additional strontium-bearing phases such as celestine. For all samples from Madagascar and for a major portion of those from Brazil, in the Al2O3 versus P2O5 plot a positive correlation can be seen (Fig. 3C). The good correlation is produced by the roughly similar P2O5 to Al2O3 ratio in the aluminum phosphates and the absence of other aluminum-bearing minerals in the rock. The presence of a major amount of aluminum-rich minerals, such as the aluminosilicates and micas in some of the samples, produces the cluster along the abscissa. In the Al2O3 versus Sr plot, the samples from Madagascar show a very steep positive correlation (Fig. 3D). This correlation is due to the substitution of strontium for calcium in the goyazite-crandallite minerals and the similar ratio of strontium to aluminum in the aluminum phosphates. The scatter in the samples from Brazil reflects the presence of other strontium-bearing minerals such as celestine. The Fe2O3 versus P2O5 diagram shows a very weak positive correlation, with a strong scatter (Fig. 3E). A correlation of iron and phosphorus can be deduced from field observations, too (Morteani and Ackermand, 2004). The Y versus P2O5 diagram shows a positive correlation (Fig. 3F). The correlation suggests that in the APB metaquartzites, most yttrium is bound to phosphate minerals, such as crandallite and apatite (Altschuler, 1980). Xenotime is a second potential yttrium source, but only a minor constituent of the APB metaquartzites.
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Rare Earth Element Whole-Rock Analyses Whole-rock rare earth element (REE) distribution patterns, normalized to the North American Shale Composite (NASC) (McLennan, 1989), that are typical for APB metaquartzites are given in Figure 4. As shown by Morteani and Ackermand (2001), the predominant carriers of the REE in the APB metaquartzites are zircon, monazite, and xenotime. Minor REE content is often shown by crandallite. The shape of the whole-rock REE distribution pattern is produced predominantly by differences in the respective amounts of zircon, monazite, and xenotime. Zircon and xenotime preferentially accommodate the heavy REE, but monazite prefers the light REE. All three minerals may display negative europium and cerium anomalies reflecting their provenance and later sedimentary and metamorphic crystallization.
10
1
0.1
0 La Ce Pr Nd
Mineral Chemistry Representative electron microprobe point analyses of phosphate-bearing minerals from the APB metaquartzites are given in Table 2. Lazulite, MgAl2(PO4)2(OH)2, and goyazite-crandallite are the most common phosphate-bearing minerals in the APB metaquartzites. Microprobe analyses show solid solution toward scorzalite Fe2+ Al2(PO4)2(OH)2. Augelite, Al2(PO4)(OH)3, deviates only slightly from the ideal composition. Berlinite, AlPO4, has optical properties that are nearly identical to those of quartz. Therefore, a reliable identification requires microprobe analysis. Trolleite, Al4(PO4)3(OH)3, is often found associated with augelite at the grain boundary between goyazite and lazulite as well as between berlinite and augelite. Minerals with a composition between the end-members svanbergite, SrAl3(PO4)(SO4)(OH)6; goyazite, SrAl3(PO4)(HPO4)(OH)6; and crandallite, CaAl3(PO4)(HPO4)(OH)6, are the main phosphate carriers in the metaquartzites of all studied localities (Fig. 5). In all these minerals, the woodhouseite component is greatly subordinate, and woodhouseite is lacking. Amblygonite has a composition close to the theoretical formula (Li, Na)(Al,Fe)PO4(F,OH), with Fe2O3 contents of up to 0.2 wt%. Wavellite, Al3(PO4)2(OH,F)3 5H2O, was recognized only by microprobe analysis. It is found in some samples in apparently stable association with berlinite, amblygonite, and trolleite. Apatite has a rather homogeneous chemical composition, with maximum FeO and SrO contents of 1.1 and 1.8 wt%, respectively. Celestine, anhydrite, tourmaline, dumortierite, kyanite, andalusite, muscovite, and hematite are also major constituents of a few samples. Rutile, zircon, monazite, and xenotime are minor constituents, and in places are concentrated in strings with hematite and tourmaline. Their mineralogy and texture suggest that these strings are black sand levels marking the original sedimentary stratification of the protolith. Metamorphic foliation commonly cuts across these strings.
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Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
Figure 4. Whole-rock rare earth element (REE) distribution patterns of the aluminum phosphate–bearing metaquartzites from Madagascar, normalized to the North American Shale Composite (NASC) (McLennan, 1989). Sample numbers are given. For further explanations, see text.
Phosphorites Whole-Rock Major and Trace Element Composition The P2O5 contents of the phosphate sediments and phosphorites listed earlier (under “Phosphorites” in the section headed “Geological Setting and Existing Data”) are between 4% and 30%, in most cases ranging from 10% to 20%. Due to their apatite mineralogy, the phosphate sediment components show a positive correlation with calcium, strontium, and fluorine, but there is no correlation between phosphorus and aluminum or iron (McClellan and Saavedra, 1986). With local variations, the trace elements Ag, Cd, La, Mo, Pb, Se, Sr, U, Y, Yb, and Zn are commonly enriched in phosphate sediments compared to shale, and B, Co, Ga, Hg, Li, Sn, Ti, and Zr are depleted (Altschuler, 1980). Rare Earth Elements REE are enriched in phosphorites by a factor of 2 to 6 compared to shale (Altschuler, 1980). The shale-normalized REE distributions of many phosphorites show a distinct negative cerium anomaly (−0.5 to −0.8), but flat shalelike patterns, positive cerium anomalies, and middle REE enrichment have also been found (Jarvis et al., 1994). The cerium anomaly might in many cases be determined by the oxidation level during apatite precipitation; e.g., a negative anomaly indicates well-oxygenated conditions during precipitation of the calcium phosphates. Mineral Chemistry All phosphorites and phosphate sediments from the deposits listed earlier are composed of carbonate fluorapatite (francolite) or fluorapatite. Francolite is the primary mineral in all sedimentary phosphate deposits. Burial diagenesis, metamorphism, and weathering cause a loss of CO2 from the lattice and alter the
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TABLE 2. REPRESENTATIVE ELECTRON MICROPROBE POINT ANALYSES OF SELECTED PHOSPHATE MINERALS FROM ALUMINUM PHOSPHATE–BEARING METAQUARTZITES FROM MADAGASCAR AND BRAZIL Phase Ber Wav Trol Apa Laz Aug Amb Goy Sample no. 10836 10836 10836 10836 10826 10835 10836 10836 10836 10826 10836 Percent SiO2 0.05 0.01 0.03 b.d.l. b.d.l. 0.09 0.02 0.07 0.06 b.d.l. b.d.l. TiO2 0.06 b.d.l. b.d.l. 0.05 b.d.l. 0.32 0.17 b.d.l. b.d.l. 0.31 0.15 Al2O3 41.20 41.82 36.55 44.79 b.d.l. 31.08 30.16 50.15 32.96 30.95 26.45 Cr2O3 0.03 b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. 0.01 b.d.l. b.d.l. b.d.l. Fe2O3 4.88 0.13 1.45 9.23 FeO 0.17 0.06 1.12 1.11 4.74 9.28 0.18 MnO 0.00 b.d.l. 0.03 b.d.l. 0.23 b.d.l. 0.13 b.d.l. b.d.l. 0.02 0.03 NiO b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. CoO b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. MgO b.d.l. 0.01 0.06 0.01 0.13 11.22 8.01 0.06 0.05 b.d.l. 0.02 ZnO b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. CaO b.d.l. 0.02 0.01 0.03 53.52 b.d.l. b.d.l. b.d.l. 0.01 2.59 2.17 BaO b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. SrO b.d.l. b.d.l. 0.02 b.d.l. 1.75 b.d.l. b.d.l. b.d.l. 0.03 17.27 15.17 Na2O 0.02 0.01 b.d.l. b.d.l. 0.08 b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. 0.33 K2O b.d.l. 0.01 0.01 b.d.l. 0.01 0.01 b.d.l. 0.02 b.d.l. 0.01 0.02 P2O5 58.67 57.98 37.34 50.87 42.72 48.28 45.72 37.22 49.61 32.74 31.76 SO3 0.01 b.d.l. b.d.l. 0.01 0.04 b.d.l. 0.05 0.03 0.04 0.53 0.89 Total 100.21 99.92 78.93 96.88 99.88 95.74 93.54 87.69 82.94 85.87 86.22 Stoichiometry Si Ti Al Cr 3+ Fe 2+ Fe Mn Ni Co Mg Zn Ca Ba Sr Na K P S Total cations
P=1
P=2
0.001 0.001 0.979
1.004
0.003
0.001
P=3
0.002 2.726
P=6 0.001
0.003 3.680
P=4 0.009 0.024 3.585
0.002 0.013 3.672
0.154 0.032
0.388
0.802 0.011
1.637
0.001
0.232 0.065 0.002
P=4
0.009
0.006
7.499 0.001 0.012
3.697
P+S=4
0.033 5.210
0.017 4.559
0.156
1.016
0.002
0.004
0.014
0.006
0.001
0.032
0.001
0.002
9.540
0.001
0.396
0.340
0.168 0.025 0.002 6.018 0.004 16.084
0.002
1.430
3.997 0.002 7.726
0.002 3.959 0.041 11.230
1.287 0.094 0.004 3.933 0.070 11.326
0.001 0.001 1.001
1.000
0.001 2.000
3.002
1.987
2.007
4.969
6.754
0.001 4.001 9.645
1.234
P=4
3.999 0.003 9.735
0.011
0.003 3.998 0.002 11.536
0.007
0.004
Oxygens 3.98 4.01 9.45 13.10 25 17.47 17.59 21.30 15.58 19.97 20.12 XMg 0.15 0.81 0.60 P + S (calc) 1.00 1.00 2.00 3.00 4.00 4.00 4.00 4.00 4.00 4.00 Notes: Ber—berlinite; Wav—wavellite; Trol—trolleite; Apa—apatite; Laz—lazulite; Aug—augelite; Amb—amblygonite; Goy—goyazite; b.d.l.—below detection limit in routine analysis. Fe total given as FeO, calculated as Fe2+/Fe3+; goyacite without Ba; amblygonite without Li.
Aluminum phosphate in Proterozoic metaquartzites goyazite SrAl 3(PO 4 ) (HPO4)(OH) 6
svanbergite SrAl3(PO 4)(SO 4 )(OH) 6 1.0 Brazil
0.6
Brazil
0.4
crandallite group
Madagascar
S r/(Sr+C a)
beudantite group
0.8
0.2
0.0 0.0
0.2
woodhouseite CaAl3 (PO 4 )(SO4)(OH)6
0.4
0.6
(P-1)/[(P-1)+S]
0.8
1.0
crandallite CaAl3(PO 4 ) (HPO4 ) (OH)6
Figure 5. Composition of the quaternary phosphates svanabergite, goyazite, crandallite, and woodhouseite from the aluminum phosphate–bearing metaquartzites of Madagascar and Brazil (Morteani and Ackermand, 2001, 2004, 2006).
composition of francolite toward fluorapatite (McClellan and Saavedra, 1986). Aluminum phosphate minerals wavellite, millisite, crandallite, lacroixite, and variscite occur as secondary phosphate minerals (McClellan and Saavedra, 1986). Goyazite was reported from the deposits in the Thelon basin. The fluor hydroxyapatite described from the Khubsugul basin is considered of secondary origin. DISCUSSION APB Metaquartzites Metamorphic History The actual mineral associations and textures in the APB metaquartzites studied here reflect the pressure and temperature (P-T) conditions of the Pan-African or Brasiliano tectonothermal event (750–500 Ma). A definite estimate of the P-T conditions is hampered by incomplete experimental data on the stability of the different phosphates and by ambiguities in the interpretation of the microtextures produced by the polystage Pan-African or Brasiliano tectonothermal event. Summarizing the results of Morteani and Ackermand (2001, 2004, 2006), the P-T conditions in the three given localities are upper greenschist to lower amphibolite facies. Uncertainty about the mineral reactions induced by metamorphism makes it difficult to define the phosphate paragenesis of the protolith of the APB metaquartzites. From the mineral reactions given by Morteani and Ackermand (2006); the aluminum phosphates described, e.g., by Rasmussen (1996)
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and Spötl (1990), from unmetamorphosed sandstones; and textural considerations regarding the given APB metaquartzites, it appears that the authigenic mineral paragenesis was very likely lazulite, augelite, amblygonite, xenotime, and minerals with a composition intermediate between goyazite, svanbergite, and crandallite. Minerals with a gorceixite composition, as described by Rasmussen (1996), could not be found. A careful study of the phosphate minerals in unmetamorphosed marine sandstones deposited along the Gondwana margin and in the Rheic Ocean and additional experimental data on the stability of the Al, Fe, Sr, and Ca phosphates are needed in the future. All present goyazite analyses, plotted in Figure 5, show a composition between goyazite and crandallite and between goyazite and svanbergite. The analyses support the existence of a binary continuous solid solution between goyazite and crandallite (crandallite group), as already supposed by Wise (1975) and Morteani and Ackermand (1996), and between goyazite and svanbergite. No phosphates with a woodhouseite composition or a very marked woodhouseite component could be found. The existence of a solid solution series between svanbergite and woodhouseite (beudantite group) and between woodhouseite and crandallite remains an open question. Genesis of the Protolith Large amounts of phosphate sediments or phosphorites are unknown from the time before the Late Proterozoic (Veizer, 1988) except the unique stromatolitic occurrences in India (see “Phosphorites” in the section headed “Geological Setting and Existing Data”). The rather rare phosphate sediments older than the Neoproterozoic are commonly nonpelletal and include phosphatic bindstones or microsphorite (phosphate mudstone) lenses or concretions (e.g., see Sheldon, 1981). The textures of Neoproterozoic phosphorites resemble those of Phanerozoic deposits. The latter are composed mainly of the apatite precursor mineral francolite. The precipitation of phosphate minerals in sediments is considered a very early diagenetic process within the top few tens of centimeters of the sediment column (Glenn et al., 1994; Trappe, 1998). The studied metaquartzites show no evidence of discrete phosphate grains or biogenic structures even in samples where sedimentary structures, such as bedding, wave ripples, and so on, are preserved. A chemical precipitation of aluminum phosphates from porewater near the water-sediment interface without a direct relationship to biogenic activity is likely. The textural composition of the metaquartzites indicates diagenetic authigenesis of the phosphate minerals as cement or replacement of earlier mineral phases. A secondary nature of the aluminum phosphates as reported from weathered horizons within phosphate deposits can be excluded due to the lack of abundant calcium phosphate precursor grains. Depositional Environment Cross-bedding and ripples demonstrate sedimentation of the protolith of APB metaquartzites in an agitated environment such as the peritidal zone in the marine realm or in a fluvial to alluvial continental setting. The association with metapelites documents a
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spatial interrelation with low-energy environments either on tidal flats and in lagoons or below the wave base. A peritidal realm is supported in particular by metamorphic carbonate bindstones or stromatolites, as found within the Itremo Group in Madagascar. Water energy controlled the distribution of clay-bearing sands and pelites. The shale normalized REE pattern (Fig. 4) with a negative cerium anomaly is generally considered evidence of marine and oxygenating conditions within the water column and at the sediment surface, which does not exclude oxygendeficient conditions in the underlying soft sediment (McArthur and Walsh, 1984; Shields and Stille, 2001; Chen et al., 2003). The negative europium anomaly is characteristic of open marine seawater (Elderfield and Greaves, 1982). Fluvial supply is possible, but proof is lacking. The sedimentological and geochemical data characterize a shallow marine, partly agitated environment with oxygenated surface conditions, such as a coastal area with beaches and supratidal flats spreading seaward to the zone below the wave base on a passive margin. The zone of maximum phosphate mineral formation is the open shallow marine realm of the shelf or adjacent epicontinental basins, though phosphogenesis in peritidal environments was recognized as well (for summary see Trappe, 1998). Upwelling contributes to phosphogenesis but is not a crucial factor. Phosphorites Metamorphic History The Precambrian phosphate sediments and phosphorites consist of apatite minerals, which are stable during burial diagenesis and metamorphism. However, the primary mineral phase carbonate fluorapatite, precipitated in sediments, is CO2-rich francolite (carbonate fluorapatite), a metastable apatite modification. During thermal overprint, a loss of CO2 alters apatite composition toward fluorapatite. The alteration proceeded to a different degree in the deposits listed earlier due to their individual tectonic history. Lithofacies of the Phosphate Sediments and Phosphorites The Precambrian phosphorites and phosphate sediments differ broadly in their textural appearance. The fabrics can be divided into pristine (in situ) fabrics and granular fabrics. Pristine fabrics comprise phosphatic mudstone layers, lenses, nodules, and microconcretions. These fabrics originate from early diagenetic, cryptocrystalline precipitation of apatite within the soft sediment body (Trappe, 1998, 2001). In carbonate suites, calcareous sediment or biogenic structures are impregnated by cryptocrystalline apatite. Common phosphate intraclasts provide evidence of a very early, synsedimentary apatite precipitation. Phosphate microconcretions are a common source of phosphate pellets and are produced by winnowing and redeposition during the soft sediment stage of the host sediment (Trappe, 2001). Phosphate sediments occur within pelitic, siliceous, and calcareous marine deposystems. Granular phosphorites may be also associated with siliciclastics. From the Scandinavian deposits an association with tuffs is evident.
Depositional Environment All known Precambrian phosphate sediments and phosphorites are reported from marine sediments. The lithofacies of the phosphate-bearing successions indicates shallow marine shelf environments or epicontinental basins. Phosphogenic deposystems are a rare temporal and spatial phenomenon, but can develop in almost the entire shallow marine realm, from outer shelf to neritic environments, with a preference for temperate waters (Trappe, 1998). The previously listed Precambrian deposits cover the complete range of potential settings. Nevertheless, phosphate sediments from the time before the Neoproterozoic remain extremely rare. Ilyin (1990) plotted the Asian occurrences on a Precambrian terrane map and confirmed the shelf nature of the deposits. This correlates with results from Phanerozoic deposits (Trappe, 1994, 1998). PHOSPHOGENESIS The processes of phosphate fixation in Phanerozoic and modern sediments have been intensively investigated in the last decades, and the principal mechanisms have been revealed. The most common and best-studied mineral phase is the calcium phosphate mineral francolite, a carbonate fluorapatite (Tribble et al., 1995). Due to the very low phosphorus concentration in modern seawater (~0.6 ppm) (Krauskopf, 1977), apatite minerals do not precipitate directly from seawater. In Phanerozoic sediments, phosphate minerals are of authigenic, detrital, and biogenic (skeletal and microbial) origin; the source of phosphorus is predominantly biogenic (Föllmi, 1996; Trappe, 1998; Filippelli, 2002). Apatite minerals are precipitated in or at the sediment-water interface after mobilization of phosphorus through bacterial degradation from organic matter including fish debris or from inorganic supply of phosphate through iron or manganese oxyhydroxides. The concentration of dissolved phosphate in marine porewater below the sediment-water interface is substantially higher in comparison to that in seawater (Jarvis et al., 1994). Froelich et al. (1988), and later Heggie et al. (1990) and O’Brien et al. (1990), demonstrated the importance of iron for the phosphogenic processes. Their iron-pumping model provides a mechanism for the transport of phosphorus from the zone of maximum degradation of organic matter located in the upper soft sediment body into the zone of maximum dissolved phosphorus a few centimeters deeper in sediment. The large binding capacity of iron oxyhydroxides for phosphorus was reported, e.g., by Smeck (1985) and Filippelli (2002), and is demonstrated in the present case by the correlation of P2O5 and Fe2O3 (Fig. 3E). The kinetic regime of fluorapatite precipitation was investigated under simulated marine conditions by Van Cappellen (1991) and by Van Cappellen and Berner (1991). Their studies could identify different pathways of precipitation. Depending on the degree of supersaturation, fluorapatite is precipitated via different metastable calcium phosphate precursor phases. At a low degree of supersaturation, fluorapatite is precipitated directly from solution over a much longer induction time. Ruttenberg and Berner (1993)
Aluminum phosphate in Proterozoic metaquartzites showed evidence of a direct nucleation of apatite in the form of “dispersed phosphate precipitation” in coastal, nonupwelling sediments. The time frame of phosphogenesis varies from a scale of months to millennia. The complex pathways of phosphorus supply, concentration, and precipitation of calcium phosphate are variable and adapted to different depositional environments. Trappe (1998) summarized the various processes in a four-stage model of phosphogenesis. Today aluminum phosphates are common weathering products of phosphatic sediments (McClellan and Saavedra, 1986). Rasmussen (1996), however, emphasized the early diagenetic precipitation of REE aluminum phosphates in marine clay-bearing sands. The phosphorus source for aluminum phosphate and calcium phosphate minerals in marine sediments is the same. Consequently, supply, burial, and dissolved phosphate enrichment in the porewater must follow the pathways described earlier for both mineral groups. Major sources of aluminum in marine sediments are clay minerals. Rasmussen (1996) presented a model for the precipitation of the various aluminum phosphates. Detrital clay minerals on the surface of quartz grains adsorb phosphorus and REE from the porewater. Later during clay transformation, these elements are released again, together with aluminum from the clay mineral, allowing the local precipitation of aluminum phosphates on quartz grains. Another host for aluminum phosphate precipitation is the partial dissolution and alteration of detrital aluminosilicates. Rasmussen (1996) attributed the processes of aluminum phosphate precipitation to the sulfate reduction and methanogenesis zones due to the presence of pyrite and siderite in his studied sediments in contrast to the suboxic conditions found during apatite precipitation. The iron mineral in the Itremo Group is hematite. The mechanism is supported by the close association of clay minerals with aluminum phosphates in marine sandstones (Rasmussen, 1996) and of aluminum phosphates with muscovite and kyanite in the present case. The studies of apatite and aluminum phosphates in sediments demonstrate the concomitant formation of the minerals in the Phanerozoic and late Neoproterozoic sediments, with the geochemical characteristics of the sediments giving preference to calcium phosphate or aluminum phosphate precipitation (Rasmussen, 1996). The textural similarities and the abundance of late Neoproterozoic or Ediacaran phosphate sediments compared to Phanerozoic and modern deposits suggest similar conditions and processes of phosphate mineral formation. A marked change is indicated by the rarity of calcium phosphate sediments before the Ediacaran, whereas aluminum phosphate deposition was apparently not affected (Rasmussen, 1996; Morteani and Ackermand, 2001, 2004). Furthermore, decreasing organic burial before the Neoproterozoic (Des Marais et al., 1992) must have reduced the potential biogenic phosphorus sink into the sediment and must have emphasized the inorganic sink through the iron and manganese sorption or through loose sorption (Föllmi, 1996). Iron was derived from the oxygen-poor Proterozoic deep seawater or from reduced
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subaqueous volcanic exhalations that got their iron contents by interacting with clastic sediments (Holland, 1973, 1984; Drever, 1974) or mafic tuffs (Simonson, 1985; Jacobsen and PimentelKlose, 1988; Derry and Jacobsen, 1990). Consequently, mechanisms of phosphorus transfer into the sediment existed during early Precambrian times, allowing phosphate mineral precipitation. In modern marine environments, ~95% of the reactive phosphorus is reused by the marine ecosystem. With a decreasing amount of biomass toward the early Precambrian, the dissolved phosphorus content in the seawater could have been higher. PROTEROZOIC PHOSPHORUS CYCLE AND LIFE In the modern biosphere, phosphorus is a deficit element and controls bioproduction. Some 95% of the reactive phosphorus is rapidly exchanged between organic matter and water or soil. The sink and burial of phosphorus are mainly bound to (1) organic deposition, (2) carbonate fluorapatite mineralization, (3) iron adsorption, and (4) carbonate deposition (Ruttenberg, 1993; Föllmi, 1996; Filippelli, 2002). Before the Ediacaran, two marked changes affected the phosphorus sink in the oceans. First, large-scale calcium phosphate deposition had almost completely ceased, whereas aluminum phosphate deposition had apparently been persistent. Second, with decreasing organic deposition toward the Early Proterozoic, the rate of organic sink had decreased (Des Marais et al., 1992). The inorganic phosphorus sink should have persisted. Due to the latter, the Proterozoic seas are considered to have been characterized by relatively high concentrations of phosphorus (Einsele, 1992) or by “excess phosphate,” phosphate that is not used in the biocycle. With the evident absence of major calcium phosphate sediments and significant deposition of organic matter before the Neoproterozoic, two major mechanisms for phosphorus sink and burial are lacking. This was thought to have caused the rapid rise of phosphorus levels in seawater during the Archean and Proterozoic. The buildup of the dissolved phosphorus content in the seawater would have been crucial for the development of life and the production of biomass, thus for the formation of atmospheric oxygen during the Proterozoic. However, this study as well as the works of Rasmussen (1996) and of Morteani and Ackermand (2001, 2004) emphasizes the widespread formation of aluminum phosphate minerals during the Proterozoic and Archean. On a broader scale, the widespread occurrence of aluminum phosphates in marine sandstones, as shown by Rasmussen (1996), and the remarkably high P-T stability of aluminum phosphates suggest that aluminum phosphates should be present, but not identified, in many metaquartzites of marine origin up to the amphibolite facies (Wise and Loh, 1976; Spötl, 1990; Rasmussen, 1996; Schmid-Beurmann, 1997; Schmid-Beurmann et al., 1997, 2000; Rasmussen et al., 1998; Morteani and Ackermand, 2004). If so, aluminum phosphate precipitation reveals an additional important mechanism of phosphorus burial beyond organic deposition and strata-bound carbonate fluorapatite formation.
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Rasmussen (1996) emphasized the importance of aluminum phosphates for the modern phosphorus cycle. He calculated a burial flux of oceanic reactive phosphorus from aluminum phosphates of 7.55 × 1010 moles/year in the modern ocean. The flux nearly equals the sink from apatite (9.1 × 1010 moles/year) or from organic iron-bound and loosely adsorbed phosphorus (9.4 × 1010 moles/year). The timing and rate of the Proterozoic oceanic buildup of the dissolved phosphorus content to modern levels from weathering of the evolving continents is unknown. The proposed high oceanic phosphorus levels, or even “excess phosphate” in the Precambrian due to the formerly postulated lack of a major sink, has to be reassessed after the identification of aluminum phosphates as a major type of phosphorus burial. This additional inorganic mechanism of phosphorus adsorption onto iron complexes, redox shuttling into the sediment, and adsorption of phosphorus onto clay minerals with aluminum phosphates resulting is independent of organic sink in the Proterozoic ocean. The inorganic process reveals a source of important phosphorus burial before the Neoproterozoic. Phosphorus burial by early diagenetic aluminum phosphate precipitation indicates significantly lower phosphorus contents in the early oceans and a slower buildup of the phosphorus concentration to the threshold necessary for the development of life beyond microbial organisms. The other important phenomenon is the rarity of calcium phosphate deposition before the Ediacaran. The phenomenon could reflect a general change in ocean chemistry. Kempe and Degens (1985) postulated for the Archean to the middle Neoproterozoic a “soda ocean” with a pH of between 9 and 11, in which bicarbonate was the dominant anion, even exceeding chloride. In a generally calciumpoor, soda-rich ocean, limited calcium availability in potentially phosphogenic environments had inhibited extensive precipitation of calcium phosphate. Exceptional precipitation of calcium phosphates was related to local calcium-rich waters in areas of local oversaturation with respect to calcium carbonate (Kempe and Degens, 1985). With the transformation of the ocean into a “calcium ocean” in the Neoproterozoic in the context of enhanced organic phosphate sink, higher calcium concentrations favored the formation of widespread calcium phosphate deposits besides aluminum phosphates. Significant changes in seawater during the Precambrian were denied by others (e.g., Holland 1984). The REE studies of Bau and Dulski (1996) disagree with the concept of a “soda ocean.” Grotzinger and Kasting (1993), however, proposed a high ratio of bicarbonate to calcium for early Precambrian seawaters. During progressive evaporation of seawater, calcium would have been exhausted even before gypsum could precipitate. This abiotic surplus carbonate precipitation was extensive in the Archean and Paleoproterozoic, and only to a lesser extent than predicted by Kempe and Degens (1985) in the Middle and Neoproterozoic. The pH in the model of Grotzinger and Kasting (1993) may have been similar to the modern value of 8.1. Their model would explain the absence of larger calcium phosphate deposits in the Archean and the Paleoproterozoic, but does not explain the changes at the beginning of the Ediacaran.
The increase of the calcium level in the ocean water during the Middle and Neoproterozoic episode coincided with a drastic increase of organic carbon burial (Des Marais et al., 1992) during the Neoproterozoic and therefore with an enhanced flux of phosphorus into the sediment. In addition, sediment properties changed through the invention of bioturbation toward the Ediacaran. The commencement of large-scale calcium phosphate deposition seems more likely to have occurred as a multiprocess event rather than one controlled by a single phenomenon. CONCLUSIONS Aluminum phosphates occur as minute pore fillings or grain-lining crystals of early diagenetic origin in the investigated Proterozoic quartzites from Brazil and Madagascar and have previously been recognized in other Precambrian sandstones. The aluminum phosphates originated from the interaction of inorganic dissolved phosphorus, iron oxyhydroxy cycling, and clay minerals near the sediment surface. The resulting minerals show a remarkably high P-T stability during subsequent metamorphism. Calcium phosphate deposition remained a rather insignificant phenomenon for phosphorus fixation in sediments before the Neoproterozoic and rapidly commenced in the Ediacaran. The phenomenon is linked to the increase of organic matter deposition and the resulting flux of phosphorus into the sediment, increasing calcium levels and the reorganization of sediment properties due to bioturbation. The volumetrically minor but widespread aluminum phosphate components are considered to contribute a significant portion to the global phosphorus sink. The inorganic phosphorus burial via aluminum phosphates was particularly relevant in the Archean to the Neoproterozoic, before the beginning of widespread calcium phosphate deposition and organic matter-bound phosphorus burial in the Ediacaran. Phosphorus is a crucial nutrient for all forms of life, and therefore it is a limiting factor for bioproduction. The broad continental shelves along the northern margin of Gondwana and the developing Rheic Ocean that hosted environments for the evolution of metazoan life also provided very widespread siliciclastic deposystems for phosphorus storage via early diagenetic aluminum phosphate precipitation. Before the Ediacaran, phosphorus burial in the oceans was dominated by the aluminum phosphates instead of the calcium phosphates (apatite). The removal of phosphorus from early ocean water by aluminum phosphates suggests lower phosphorus nutrient levels than have been proposed before and a slower buildup of the oceanic phosphorus content to a threshold for the development of higher forms of life during the Proterozoic. ACKNOWLEDGMENTS We thank Christine Preinfalk (Munich) for critical reading and help with layout. The manuscript benefited greatly from the constructive comments of the reviewers, G. Shields and E. S. Grew.
Aluminum phosphate in Proterozoic metaquartzites REFERENCES CITED Almeida Abreu, P.A., 1993, A evolução geodinamica da Serra do Espinhaço meridional, Minas Gerais, Brasil [Ph.D. thesis]: Freiburg im Breisgau, Germany, Albert-Ludwigs-Universität, 150 p. Altschuler, Z.S., 1980, The geochemistry of trace elements in marine phosphorites, part I: Characteristic abundance and enrichment: Tulsa, Oklahoma, Society of Economic Paleontologists and Mineralogists Special Publication 29, p. 19–30. Banerjee, D.M., 1986, Proterozoic and Cambrian phosphorites—Regional review: Indian subcontinent, in Cook, P.J., and Shergold, J.H., eds., Phosphate deposits of the world, v. 1: Proterozoic and Cambrian phosphorites: Cambridge, England, Cambridge University Press, p. 70–90. Banerjee, D.M., Schidlowski, M., and Arneth, J.D., 1986, Genesis of Upper Proterozoic–Cambrian phosphorite deposits of India: Isotopic inference from carbonate fluorapatite, carbonate and organic carbon: Precambrian Research, v. 33, p. 239–253, doi: 10.1016/0301-9268(86)90024-0. Bau, M., and Dulski, P., 1996, Distribution of yttrium and rare-earth elements in the Penge and Kuruman iron-formations, Transvaal Supergroup, South Africa: Precambrian Research, v. 79, p. 37–55, doi: 10.1016/0301-9268 (95)00087-9. Bhatia, M.R., 1983, Plate tectonics and geochemical composition of sandstones: Journal of Geology, v. 91, p. 611–627. Brito Neves, B.B., Kawashita, K., Cordani, U.G., and Delhal, J., 1979, A evoluçao geochronológica da cordilheira do Espinhaço: Dados novos e integraçao: Revista Brasileira de Geociencias, no. 9, p. 71–85. Bruni, M.A., and Schobbenhaus Filho, C., 1976, Carta geologica do Brasil ao milionesimo: Folha Brasilia (SD 23): Ministerio das Minas e Energia, Departamento Nacional da Produção Mineral, Brasilia, p. 1–162. Cathcart, J.B., 1977, Phosphorite deposits near Patos de Minas, Minas Gerais, Brazil: Transactions of the Society of Mining Engineers of the American Institute of Mining, Mettalurgical, and Petroleum Engineers (AIME), v. 264, p. 1703–1708. Chen, D.F., Dong, W.Q., Qi, L., Chen, G.Q., and Chen, X.P., 2003, Possible REE constraints on the depositional and diagenetic environment of Doushantuo Formation phosphorites containing the earliest metazoan fauna: Chemical Geology, v. 201, p. 103–118, doi: 10.1016/S0009-2541(03)00235-3. Christie, R.L., and Sheldon, R.P., 1986, Proterozoic and Cambrian phosphorites—Regional review: North America, in Cook, P.J., and Shergold, J.H., eds., Phosphate deposits of the world, v. 1: Proterozoic and Cambrian phosphorites: Cambridge, England, Cambridge University Press, p. 101–107. Collins, A.S., Johnson, S., Fitzsimons, C.W., Powell, C.M., Hulschner, B., Abello, J., and Razakamanana, T., 2003, Neoproterozoic deformation in central Madagascar: A structural section through part of the East African Orogen, in Yoshida, M., et al., eds., Proterozoic East Gondwana: Supercontinent assembly and breakup: Geological Society of London Special Publication 296, p. 363–379. Cook, P.J., and Shergold, J.H., 1986a, Proterozoic and Cambrian phosphorites—An introduction, in Cook, P.J., and Shergold, J.H., eds., Phosphate deposits of the world, v. 1: Proterozoic and Cambrian phosphorites: Cambridge, England, Cambridge University Press, p. 1–8. Cook, P.J., and Shergold, J.H., 1986b, Proterozoic and Cambrian phosphorites—Nature and origin, in Cook, P.J., and Shergold, J.H., eds., Phosphate deposits of the world, v. 1: Proterozoic and Cambrian phosphorites: Cambridge, England, Cambridge University Press, p. 369–386. Cordani, U.G., Iyer, S.S., Taylor, P.N., Kawashita, K., Sato, K., and McReath, I., 1992, Pb-Pb, Rb-Sr, and K-Ar systematics of the Lagoa Real uranium province (south-central Bahia, Brazil) and the Espinhaço cycle (ca. 1.5– 1.0 Ga): Journal of South American Earth Sciences, v. 5, p. 33–46, doi: 10.1016/0895-9811(92)90058-7. Cox, R., Armstrong, R.A., and Ashwal, L.D., 1998, Sedimentology, geochronology and provenance of Proterozoic Itremo Group, central Madagascar, and implications for pre-Gondwana palaeogeography: Journal of the Geological Society, v. 155, p. 1009–1024. Dardenne, M.A., Trompette, R., Magalhàes, L.F., and Soares, L.A., 1986, Proterozoic and Cambrian phosphorites—Regional review: Brazil, in Cook, P.J., and Shergold, J.H., eds., Phosphate deposits of the world, v. 1: Proterozoic and Cambrian phosphorites: Cambridge, England, Cambridge University Press, p. 116–131. Derry, L.A., and Jacobsen, S.B., 1990, The chemical evolution of Precambrian seawater: Evidence from REE’s in banded iron formations: Geo-
591
chimica et Cosmochimica Acta, v. 54, p. 2965–2977, doi: 10.1016/00167037(90)90114-Z. Des Marais, D.J., Strauss, H., Summons, R.E., and Hayes, J.M., 1992, Carbon isotope evidence for the stepwise oxidation of the Proterozoic environment: Nature, v. 359, p. 605–609, doi: 10.1038/359605a0. Drever, J.I., 1974, Geochemical model for the origin of Precambrian iron formations: Geological Society of America Bulletin, v. 85, p. 1099–1106, doi: 10.1130/0016-7606(1974)85<1099:GMFTOO>2.0.CO;2. Einsele, G., 1992, Sedimentary basins: Evolution, facies and sediment budget: Berlin, Springer, 628 p. Ek, R., and Nysten, P., 1990, Phosphate mineralogy of the Hålsjöberg and Hökensås kyanite deposit: Geologiska Föreningens I Stockholm Förhandlingar, v. 112, p. 9–18. Elderfield, H., and Greaves, M.J., 1982, The rare earth elements in seawater: Nature, v. 296, p. 214–219, doi: 10.1038/296214a0. Fernandez, A., and Schreurs, G., 2003, Tectonic evolution of the Proterozoic Itremo Group metasediments in central Madagascar, in Yoshida, M., et al., eds., Proterozoic east Gondwana: Supercontinent assembly and breakup: Geological Society of London Special Publication 206, p. 318–399. Filippelli, G.M., 2002, The global phosphorus cycle, in Kohn, M.J., et al., eds., Phosphates: Geochemical, geobiological, and materials importance: Reviews in Mineralogy and Geochemistry, v. 48, p. 391–418. Föllmi, K.B., 1996, The phosphorus cycle, phosphogenesis and marine phosphaterich deposits: Earth-Science Reviews, v. 40, p. 55–124, doi: 10.1016/00128252(95)00049-6. Froelich, P.N., Arthur, M.A., Burnett, W.C., Deakin, M., Hensley, V., Jahnke, R., Kaul, L., Kim, K.-H., Roe, K., Soutar, A., and Vathakanon, C., 1988, Early diagenesis of organic matter in Peru continental margin sediments: Phosphorite precipitation: Marine Geology, v. 80, p. 309–349, doi: 10.1016/ 0025-3227(88)90095-3. Gabaldón López, V., Hernández Urroz, J., Lorenzo Alvarez, S., Picart Boira, J., Santamará Casanova, J., and Solé Pont, F.J., 1989, Sedimentary facies and stratigraphy of Precambrian–Cambrian phosphorites on the Valdelacasa Anticline, Central Iberian, in Notholt, A.J.G., et al., eds., Phosphate deposits of the world, v. 2: Phosphate rock resources: Cambridge, England, Cambridge University Press, p. 422–428. Glenn, C.R., Föllmi, K.B., Riggs, S.R., Baturin, G.N., Grimm, K.A., Trappe, J., Abed, A.M., Galli-Oliver, C., Garrison, R.E., Ilyin, A.V., Jehl, C., Roehrlich, V., Sadaqah, R.M.Y., Schidlowski, M., Sheldon, R.E., and Siegmund, H., 1994, Phosphorus and phosphorites: Sedimentology and environments of formation: Eclogae Geologicae Helvetiae, v. 87, no. 3, p. 747–788. Golonka, J., Ross, M.I., and Scotese, C.R., 1994, Phanerozoic and paleogeographic and paleoclimatic modelling maps, in Beauchamp, B., et al., eds., Carboniferous to Jurassic Pangea: Canadian Society of Petroleum Geologists Memoir 17, p. 1–47. Grotzinger, J.P., and Kasting, J.F., 1993, New constraints on Precambrian ocean composition: Journal of Geology, v. 101, p. 235–243. Heggie, D.T., Skyring, G.W., O’Brien, G.W., Reimers, C., Herczeg, A., Moriarty, D.J.W., Burnett, W.C., and Milnes, A.R., 1990, Organic carbon cycling and modern phosphorite formation on the East Australian continental margin: An overview: Geological Society of London Special Publication 52, p. 87–117. Holland, H.D., 1973, The oceans: A possible source of iron in iron-formations: Economic Geology and the Bulletin of the Society of Economic Geologists, v. 68, p. 1169–1172. Holland, H.D., 1984, The chemical evolution of the atmosphere and oceans: Princeton, New Jersey, Princeton University Press, 598 p. Howard, P.F., 1986, Proterozoic and Cambrian phosphorites—Regional review: Australia, in Cook, P.J., and Shergold, J.H., eds., Phosphate deposits of the world, v. 1: Proterozoic and Cambrian phosphorites: Cambridge, England, Cambridge University Press, p. 20–41. Ilyin, A.V., 1990, Proterozoic supercontinent, its latest Precambrian rifting, breakup, dispersal into smaller continents, and subsidence of their margins: Evidence from Asia: Geology, v. 18, p. 1231–1234, doi: 10.1130/00917613(1990)018<1231:PSILPR>2.3.CO;2. Ilyin, A.V., Ochir, L., and Ratnikova, G.I., 1989, The Khubsgul phosphorite deposit, Mongolian People’s Republic, in Notholt, A.J.G., et al., eds., Phosphate deposits of the world, v. 2: Phosphate rock resources. Cambridge, England, Cambridge University Press, p. 482–484. Jacobsen, S.B., and Pimentel-Klose, M.R., 1988, A Nd isotopic study of the Hamersley and Michipicoten banded iron formations: The source of REE and Fe in Archean oceans: Earth and Planetary Science Letters, v. 87, p. 29–44, doi: 10.1016/0012-821X(88)90062-3.
592
Morteani et al.
Jarvis, I., Burnett, W.C., Nathan, Y., Almbaydin, F.S.M., Attia, A.K.M., Castro, L.N., Flicoteaux, R., Hilmy, M.E., Husain, V., Qutawnah, A.A., Serjani, A., and Zanin, Y.N., 1994, Phosphorite geochemistry: State-of-the-art and environmental concerns: Eclogae Geologicae Helvetiae, v. 87, p. 643–700. Kalt, H., 1991, Untersuchungen zur Schertektonik in der südlichen Serra do Espinhaço und im angrenzenden Teil des São Francisco Beckens, Minas Gerais, Brasilien [Ph.D. thesis]: Freiburg im Breisgau, Germany, AlbertLudwigs-Universität, 103 p. Kempe, S., and Degens, E.T., 1985, An early soda ocean?: Chemical Geology, v. 53, p. 95–108, doi: 10.1016/0009-2541(85)90023-3. Krasilnikova, N.A., and Ilyin, A.V., 1989, Late Precambrian–Early Cambrian phosphorite province of South Siberia, in Notholt, A.J.G., et al., eds., Phosphate deposits of the world, v. 2: Phosphate rock resources: Cambridge, England, Cambridge University Press, p. 514–516. Krauskopf, K.B., 1977, Introduction to geochemistry: New York, McGrawHill, 617 p. Lacroix, A., 1922–23, Minéralogie de Madagascar, 3 vols.: Paris, Challamel. Li, Y., 1986, Proterozoic and Cambrian phosphorites—Regional review: China, in Cook, P.J., and Shergold, J.H., eds., Phosphate deposits of the world, v. 1: Proterozoic and Cambrian phosphorites: Cambridge, England, Cambridge University Press, p. 42–69. López, N.J., and De Souza, J.D., 1985, Projeto mapas metalogenéticos e de previsão de recursos minerais: Carta metalogenetica: Carta de previsão de recursos minerais: Carta de previsão para planejamento de ações governamentais: Folha SD. 23-X-B, Ibotirama, Região Nordeste: Brasilia, DNPM, scale 1:250.000. McArthur, J.M., and Walsh, J.N., 1984, Rare-earth geochemistry of phosphorites: Chemical Geology, v. 47, p. 191–220, doi: 10.1016/0009-2541(84)90126-8. McClellan, G.H., and Saavedra, F.N., 1986, Chemical and mineral characteristics of some Cambrian and Precambrian phosphorites, in Cook, P.J., and Shergold, J.H., eds., Phosphate deposits of the world, v. 1: Proterozoic and Cambrian phosphorites: Cambridge, England, Cambridge University Press, p. 244–267. McLennan, S.M., 1989, Rare earth elements in sedimentary rocks: Influence of provenance and sedimentary processes, in Lipin, B.R., and McKay, G.A., eds., Geochemistry and mineralogy of rare earth elements: Chantilly, Virginia, Mineralogical Society of America, Reviews in Mineralogy, v. 21, p. 169–200. Moine, B., 1967, Relations stratigraphiques entre la série “schiste-quartzite-calcaire” et les gneiss environants (centre ouest de Madagascar): Comptes Rendus de la Semmaine Géologique de Madagascar, Tananarive, p. 49–53. Morteani, G., and Ackermand, D., 1996, Aluminium phosphates in muscovitekyanite metaquartzite from Passo di Vizze (Alto Adige, NE Italy): European Journal of Mineralogy, v. 8, p. 853–869. Morteani, G., and Ackermand, D., 2001, Aluminium-phosphates and -borosilicates in muscovite-kyanite metaquartzites near Diamantina (Minas Gerais, Brazil): Periodico di Mineralogia, v. 70, p. 111–129. Morteani, G., and Ackermand, D., 2004, Mineralogy and geochemistry of Alphosphate and Al-borosilicate-bearing metaquartzites of the northern Serra do Espinhaço (State of Bahia, Brazil): Mineralogy and Petrology, v. 80, p. 59–81, doi: 10.1007/s00710-003-0019-z. Morteani, G., and Ackermand, D., 2006, Mineralogy, geochemistry and petrology of an amphibolite-facies aluminium-phosphate and borosilicate (APB)-bearing quartzite from the Mesoproterozoic Itremo Group (Central Madagascar): Neues Jahrbuch für Mineralogie, Abhandlungen, v. 182, p. 123–148; doi: 10.1127/0077-7757/2006/0036. Notholt, A.J.G., and Brasier, M.D., 1986, Proterozoic and Cambrian phosphorites—Regional review: Europe, in Cook, P.J., and Shergold, J.H., eds., Phosphate deposits of the world, v. 1: Proterozoic and Cambrian phosphorites: Cambridge, England, Cambridge University Press, p. 91–100. Notholt, A.J.G., Sheldon, R.P., and Davidson, D.F., 1989, Asia—Introduction, in Notholt, A.J.G., et al., eds., Phosphate deposits of the world, v. 2: Phosphate rock resources: Cambridge, England, Cambridge University Press, p. 437–442. O’Brien, G.W., Milnes, A.R., Veeh, H.H., Heggie, D.T., Riggs, S.R., Cullen, D.J., Marshall, J.F., and Cook, P.J., 1990, Sedimentation dynamics and redox iron-cycling: Controlling factors for apatite-glauconite association on the East Australian continental margin: Geological Society of London Special Publication 52, p. 442–444. Rasmussen, B., 1996, Early-diagenetic REE-phosphate minerals (florencite, gorceixite, crandallite, and xenotime) in marine sandstones: A major sink for oceanic phosphorus: American Journal of Science, v. 296, p. 601–632.
Rasmussen, B., Buick, R., and Taylor, W.R., 1998, Removal of oceanic REE by authigenic precipitation of phosphatic minerals: Earth and Planetary Science Letters, v. 164, p. 135–149, doi: 10.1016/S0012-821X(98)00199-X. Ruttenberg, K.C., 1993, Reassessment of the oceanic residence time of phosphorus: Chemical Geology, v. 107, p. 405–407, doi: 10.1016/0009-2541(93)90220-D. Ruttenberg, K.C., and Berner, R.A., 1993, Authigenic apatite formation and burial in sediments from non-upwelling, continental margin environments: Geochimica et Cosmochimica Acta, v. 57, p. 991–1007, doi: 10.1016/00167037(93)90035-U. Schmid-Beurmann, P., 1997, Phosphat-Mineralogie und ihre experimentellen materialwissenschaftlichen Aspekte dargestellt am Beispiel des Lazulith-Typs [Habilitationsschrift]: Kiel, Germany, Christian Albrechts Universität, 153 p. Schmid-Beurmann, P., Morteani, G., and Cemicˇ, L., 1997, Experimental determination of the upper stability of scorzalite, FeAl2[OH/PO4]2, and the occurrence of minerals with a composition intermediate between scorzalite and lazulite (ss) up to conditions of the amphibolite facies: Mineralogy and Petrology, v. 61, p. 211–222, doi: 10.1007/BF01172485. Schmid-Beurmann, P., Knitter, S., and Cemicˇ, L., 2000, P-T stability of scorzalite of the lazulite-scorzalite solid-solution series: Mineralogy and Petrology, v. 70, p. 55–71, doi: 10.1007/s007100070013. Sheldon, R.P., 1981, Ancient marine phosphorites: Annual Review of Earth and Planetary Sciences, v. 9, p. 251–284, doi: 10.1146/annurev.ea.09.050181.001343. Shields, G., and Stille, P., 2001, Diagenetic constraints on the use of cerium anomalies as paleoseawater redox proxies: An isotopic and REE study of Cambrian phosphorites: Chemical Geology, v. 175, p. 29–48, doi: 10.1016/S0009-2541(00)00362-4. Simonson, B., 1985, Sedimentological constraints on the origin of Precambrian iron formations: Geological Society of America Bulletin, v. 96, p. 244– 252, doi: 10.1130/0016-7606(1985)96<244:SCOTOO>2.0.CO;2. Slansky, M., 1986, Proterozoic and Cambrian phosphorites—Regional review: West Africa, in Cook, P.J., and Shergold, J.H., eds., Phosphate deposits of the world, v. 1: Proterozoic and Cambrian phosphorites: Cambridge, England, Cambridge University Press, p. 108–115. Smeck, N.E., 1985, Phosphorus dynamics in soils and landscapes: Geoderma, v. 36, p. 185–199, doi: 10.1016/0016-7061(85)90001-1. Spötl, C., 1990, Authigenic aluminium phosphates-sulphates in sandstones of the Mitterberg formation, Northern Calcareous Alps, Austria: Sedimentology, v. 37, p. 837–845, doi: 10.1111/j.1365-3091.1990.tb01828.x. Tavora, F.J., Cordani, U.G., and Kawashita, K., 1967, Determinações de idade potassio-argônio em rochas da região central da Bahia: Anais de XXI Congreso Brasileira de Geologia, Curitiba, p. 234–244. Trappe, J., 1994, Pangean phosphorites—Ordinary phosphorite genesis in an extraordinary world? in Beauchamp, B., et al., eds., Carboniferous to Jurassic Pangea: Calgary, Alberta, Canadian Society of Petroleum Geologists Memoir 17, p. 469–478. Trappe, J., 1998, Phanerozoic phosphorite depositional systems—A dynamic model for a sedimentary resource system: Lecture Notes in Earth Science, v. 74, 316 p. Trappe, J., 2001, A nomenclature system for granular phosphorites according to depositional texture: Sedimentary Geology, v. 145, p. 135–150, doi: 10.1016/S0037-0738(01)00103-8. Tribble, J.S., Arvidson, R.S., Lane, M., III, and Mackenzie, F.T., 1995, Crystal chemistry, and thermodynamic and kinetic properties of calcite, dolomite, apatite, and biogenic silica: Applications to petrologic problems: Sedimentary Geology, v. 95, p. 11–37, doi: 10.1016/0037-0738(94)00094-B. Trompette, R., 1989, Phosphorites of the northern Volta Basin (Burkina Faso, Niger and Benin, in Notholt, A.J.G., et al., eds., Phosphate deposits of the world, v. 2: Phosphate rock resources: Cambridge, England, Cambridge University Press, p. 214–218. Van Cappellen, P., 1991, The formation of marine apatite: A kinetic study [Ph.D. thesis]: New Haven, Connecticut, Yale University, 240 p. Van Cappellen, P., and Berner, R.A., 1991, Fluorapatite crystal growth from modified sea water: Geochimica et Cosmochimica Acta, v. 55, p. 1219– 1234, doi: 10.1016/0016-7037(91)90302-L. Veizer, J., 1988, The evolving exogenic cycle, in Gregor, C.N., et al., eds., Chemical cycles in the evolution of the earth: New York, Wiley, p. 175–220. Wise, W.S., 1975, Solid solution between the alunite, woodhouseite and crandallite mineral series: Neues Jahrbuch für Mineralogie, v. 12, p. 540–545. Wise, W.S., and Loh, S.E., 1976, Equilibria and origin of minerals in the system Al2O3-AlPO4-H2O: American Mineralogist, v. 61, p. 409–413. MANUSCRIPT ACCEPTED BY THE SOCIETY 3 OCTOBER 2006
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Preface 1. The assembly of West GondwanaThe view from the Rio de Ia Plata craton Kerstin Saalmann, Leo 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
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)
Ulf Linnemann, Axel Gerdes, Kerstin Drost, and Bernd Buschmann 4. The Lausitz graywackes, Saxo-Thuringia, Germany-Witness to the Cadomian orogeny
Helga Kemnitz 5. Paleontological data from the Early Cambrian of Germany and paleobiogeographical implications for the configuration of central Peri gondwana
Olaf Elicki 6. The Variscan orogeny in the Saxo-Thuringian zone-Heterogenous overprint of Cadomian/ Paleozoic Peri-Gondwana crust 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)
Christoph Breitkreuz, Allen Kennedy, Marion Geif31er, Bodo-Carlo Ehling, Jilrgen 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)
Kerstin Drost, Rolf L. Romer, Ulf Linnemann, Oldfich 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 Christian Pin, R. Kryza, T. Oberc-Dziedzic, S. Mazur, K. Turniak, and Jarmila Waldhausrova 10. Sm-Nd isotope and trace element study of Late Proterozoic meta basalts ("spilites") from the Central Barrandian domain (Bohemian Massif, Czech Republic)
Christian Pin and Jarmila Waldhausrova 11. Structural evolution ofthe Prague synform (Czech Republic) during Silurian times:
An AMS, rock magnetism, and paleomagnetic study of the SvatY Jan pod Skalou dikes. Consequences for the nappes emplacement
Tahar Ai'fa, Petr Pruner, Martin Chadima, and Petr Storch 12. Cadomian and Variscan metamorphic events in the Leon domain (Armorican Massif. France): P-T data and EMP monazite dating
Bernhard Schulz, Erwin Krenn, Fritz Finger, Helene Bratz, 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
Gabriel Gutierrez-Alonso, Javier FernandezSuarez, Juan Carlos Gutierrez-Marco, Fernando Corfu, J. Brendan Murphy, and Mercedes Suarez 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
Miguel Lopez-Plaza, Mercedes Peinado, Francisco-Javier Lopez-Mora, M. Dolores Rodrfguez-Aionso, Asuncion Carnicero, M. Piedad Franco, Juan Carlos Gonzalo, and Marina Navidad 15. Tectonic evolution of the upper allochthon of the Ordenes complex (northwestern Iberian Massif): Structural constraints to a polyorogenic peri-Gondwanan terrane
Juan Gomez Barreiro, Jose R. Martinez Catalan, Ricardo Arenas, Pedro Castifieiras, Jacobo Abati, Florentino Dfaz Garcfa, and Jan R. Wijbrans 16. Crustal growth and deformational processes in the northern Gondwana margin: Constraints from the Evora Massif (OssaMorena zone, southwest Iberia, Portugal) M. Francisco Pereira, J. Brandiio Silva,
Martim Chichorro, Patricia Moita, Jose F. Santos, Arturo Apraiz, and Cristina Ribeiro 17. The Lower-Middle Cambrian boundary in the Mediterranean subprovince
Rodolfo Gozalo, E/adio Lifian, Marfa Eugenia Dies Alvarez, Jose Antonio Gamez 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
MartinS. Oczlon, Antoneta Seghedi, and Charles W Carrigan 19. Crete and the Minoan terranes: Age constraints from U-Pb dating of detrital zircons G. Zulauf, S.S. Romano, W Dorr, 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 J. Brendan Murphy 21. Vestige of the Rheic Ocean in North America: The Acatlcln Complex of southern Mexico
R. Damian Nance, Brent V. Miller, J. Duncan Keppie, J. Brendan Murphy, and
Jaroslav Dostal 22. Provenance of the Granjeno Schist, Ciudad Victoria, Mexico: Detrital zircon U-Pb age constraints and implications for the Paleozoic paleogeography of the Rheic Ocean
R. Damian Nance, Javier Fernandez-Suarez, J. Duncan Keppie, Craig Storey, and Teresa E.
Jeffries 23. Ordovician calc-alkaline granitoids in the Acatlim Complex, southern Mexico: Geochemical and geochronologic data and implications for the tectonics of the Gondwanan margin of the Rheic Ocean Brent V. Miller, Jaroslav Dostal, J. Duncan
Keppie, R. Damian Nance, Amabel OrtegaRivera, and James K.W Lee 24. Ordovician-Devonian oceanic basalts in the Cosoltepec Formation, Acatlcln Complex, southern Mexico: Vestiges of the Rheic Ocean? J. Duncan Keppie, Jarosla v Dostal, and
Mariano Elfas-Herrera 25. P-T-t constraints on exhumation following subduction in the Rheic Ocean from eclogitic rocks in the Acatlan Complex of southern Mexico 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
Sven 0. Egenhoff 27. A Late Ordovician ice sheet in South America: Evidence from the Cancaiiiri tillites, southern Bolivia
Frank Schonian and Sven 0. Egenhoff 28. Sedimentary basins in the southwestern Siberian craton: Late Neoproterozoic-Early Cambrian rifting and collisional events
Julius Konstantinovich Sovetov, Anna Evgen'evna Kulikova, and Maxim Nikolaevich Medvedev 29. Aluminum phosphate in Proterozoic meta quartzites: Implications for the Precambrian oceanic P budget and development of life
Giulio Morteani, Dietrich Ackermand, and Jorg Trappe Index IS BN 978-0-8137-2423-2
