Late Paleozoic Glacial Events and Postglacial Transgressions in Gondwana (GSA Special Paper 468)

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Late Paleozoic Glacial Events and Postglacial Transgressions in Gondwana

edited by Oscar R. López-Gamundí Hess Corporation One Allen Center 500 Dallas Street Houston, Texas 77002 USA and Luis A. Buatois Department of Geological Sciences University of Saskatchewan 114 Science Place Saskatoon, SK S7N 5E2 Canada

Special Paper 468 3300 Penrose Place, P.O. Box 9140

Boulder, Colorado 80301-9140, USA

2010

Copyright © 2010, The Geological Society of America (GSA), Inc. 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. GSA provides this and other forums for the presentation of diverse opinions and positions by scientists worldwide, regardless of their race, citizenship, gender, religion, or political viewpoint. Opinions presented in this publication do not reflect official positions of the Society. 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 Editors: Marion E. Bickford and Donald I. Siegel Library of Congress Cataloging-in-Publication Data Late Paleozoic glacial events and postglacial transgressions in Gondwana / edited by Oscar R. LópezGamundí and Luis A. Buatois. p. cm. — (Special paper ; 468) Includes bibliographical references. ISBN 978-0-8137-2468-3 (pbk.) 1. Geology, Stratigraphic—Paleozoic. 2. Glacial landforms—Gondwana (Continent) 3. Periglacial processes—Gondwana (Continent) 4. Gondwana (Continent) I. López-Gamundí, Oscar R. II. Buatois, Luis A. QE654.L375 2010 551.7′2—dc22 2010019884 Cover, front: Map showing the distribution of glacial basins in Gondwana. From Isbell, J.L., “Environmental and paleogeographic implications of glaciotectonic deformation of glaciomarine deposits within Permian strata of the Metschel Tillite, southern Victoria Land, Antarctica” (Chapter 3, this volume). Back: Postglacial transgressive scenarios. (Upper left) Fjord environment influenced by extreme freshwater discharge from retreating glaciers. A freshwater ichnofauna occurs in the transgressive deposits (Guandacol Formation). (Lower right) Coastal environment without direct influence of freshwater influx from melting ice masses. A freshwater ichnofauna is present in coastal lakes and temporally inundated flood plains (Trace-Fossil Assemblage 1). A brackish-water ichnofauna (Trace-Fossil Assemblages 2 and 3) occurs in distal-bay facies (Tupe Formation). From Desjardins, P.R., Buatois, L.A., Mángano, M.G., and Limarino, C.O., “Ichnology of the latest Carboniferous–earliest Permian transgression in the Paganzo Basin of western Argentina: The interplay of ecology, sea-level rise, and paleogeography during postglacial times in Gondwana” (Chapter 8, this volume).

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Contents

Introduction: Late Paleozoic glacial events and postglacial transgressions in Gondwana . . . . . . . . . . v Oscar R. López-Gamundí and Luis A. Buatois 1. Transgressions related to the demise of the Late Paleozoic Ice Age: Their sequence stratigraphic context . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1 Oscar R. López-Gamundí 2. From bergs to ergs: The late Paleozoic Gondwanan glaciation and its aftermath in Saudi Arabia . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 37 John Melvin, Ronald A. Sprague, and Christian J. Heine 3. Environmental and paleogeographic implications of glaciotectonic deformation of glaciomarine deposits within Permian strata of the Metschel Tillite, southern Victoria Land, Antarctica . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 81 John L. Isbell 4. Formation of euxinic lakes during the deglaciation phase in the Early Permian of East Africa . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 101 Thomas Kreuser and Gebretinsae Woldu 5. Stratigraphic and paleofloristic record of the Lower Permian postglacial succession in the southern Brazilian Paraná Basin . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 113 Roberto Iannuzzi, Paulo A. Souza, and Michael Holz 6. “Levipustula Fauna” in central-western Argentina and its relationships with the Carboniferous glacial event in the southwestern Gondwanan margin . . . . . . . . . . . . . . . . . . . . 133 Gabriela A. Cisterna and Andrea F. Sterren 7. Ichnology of late Paleozoic postglacial transgressive deposits in Gondwana: Reconstructing salinity conditions in coastal ecosystems affected by strong meltwater discharge . . . . . . . . . . . . 149 Luis A. Buatois, Renata G. Netto, and M. Gabriela Mángano 8. Ichnology of the latest Carboniferous–earliest Permian transgression in the Paganzo Basin of western Argentina: The interplay of ecology, sea-level rise, and paleogeography during postglacial times in Gondwana . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 175 Patricio R. Desjardins, Luis A. Buatois, M. Gabriela Mángano, and Carlos O. Limarino 9. Reconstruction of a high-latitude, postglacial lake: Mackellar Formation (Permian), Transantarctic Mountains . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 193 Molly F. Miller and John L. Isbell

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The Geological Society of America Special Paper 468 2010

Introduction: Late Paleozoic glacial events and postglacial transgressions in Gondwana Oscar R. López-Gamundí Hess Corporation, 500 Dallas Street, Houston, Texas 77002, USA Luis A. Buatois Department of Geological Sciences, University of Saskatchewan, 114 Science Place, Saskatoon, SK S7N 5E2, Canada

The stratigraphic record suggests that glaciations have occurred episodically at different time intervals in Earth’s history (Crowell, 1982, 1999). One of those glaciations affected the Gondwanan Supercontinent during the late Paleozoic and constituted the longest period of continuous glaciation in the Phanerozoic (Eyles, 1993). Carboniferous to Early Permian glaciogenic successions have been known on all the subcontinents of Gondwana, most notably South America, Africa, India, and Australia, and later work expanded to Antarctica and the Middle East (Fig. 1). This glacial age can be subdivided into three distinct episodes (López-Gamundí, 1997). Glacial episodes II and III occurred during the early Late Carboniferous and the Late Carboniferous–Early Permian, respectively. An earlier, shortlived glacial episode in the Late Devonian–earliest Carboniferous (glacial episode I) identified in central and northern South America (Fig. 1) extended even further the duration of this ice age (Veevers and Powell, 1987). In general, the locus of ice cover, and its stratigraphic record, progressively moved across Gondwana from South America to Australia (Crowell, 1999), tracking the transpolar trajectory across Gondwana. This polar wander across the Gondwanan Supercontinent controlled paleolatitudes and accounts for the diachroneity of glacial episodes I, II, and III; however, the exact timing of waxing and waning of ice centers during each glacial episode (particularly for the longest-lived episode III) seemed to have been influenced by basin dynamics, topographic barriers, glaciation styles, and other local factors. The recognition of ancient glacial deposits of similar late Paleozoic age in South Africa and South America (Du Toit, 1927) helped, in conjunction with other lines of evidence, to argue in favor of the principles of seafloor spreading and indirectly to build the theory of plate tectonics (Wegener, 1915). The pioneering work during the first half of the last century was solidi-

fied subsequently by the seminal work led by John Crowell and Lawrence Frakes (Frakes and Crowell, 1967, 1969; Frakes et al., 1969; Frakes and Crowell, 1970; Crowell and Frakes, 1971a, 1971b, 1972; Frakes et al., 1971; Crowell, 1978). With the additional help of later contributions, the body of evidence about the duration, areal extent, and influence of the Late Paleozoic Ice Age (LPIA) on the biota has significantly grown. However, uncertainty remains over the exact timing of onset and demise of each glacial episode of the LPIA, particularly when attempts are made to link these glacial episodes in Gondwana with cyclothems in the Northern Hemisphere (particularly the United States and Europe), following Wanless and Shepard’s (1936) hypothesis. The picture becomes particularly blurred if, as exemplified by Wright and Vanstone (2001) for the Viséan carbonate successions in the Northern Hemisphere (UK), glacioeustatic sea-level oscillations invoked to account for high-frequency cyclicity had an approximate 100 ka periodicity, which may correspond to Milankovitch eccentricity. Thus, far field studies can sometimes be based on shaky grounds, particularly owing to the difficulty of estimating the magnitude and hierarchy of the near field glacioeustatic fluctuations (Rygel et al., 2008) and the less than optimal chronostratigraphic resolution of the Gondwanan faunas and floras associated with the LPIA. Maximum expansion of Gondwanan continental ice sheets occurred during earliest Permian time (glacial episode IV) under paleoatmospheric CO2 levels as low as the present ones to values of up to 12 times higher by the late Early Permian (Montañez et al., 2007). Widespread Early Permian (Sakmarian) collapse of ice sheets coincided with the onset of rising atmospheric CO2 levels, after which time surface temperatures and atmospheric partial pressure of carbon dioxide (pCO2) rose. The more detailed and deeper our knowledge about the LPIA gets, the better positioned

López-Gamundí, O.R., and Buatois, L.A., 2010, Introduction: Late Paleozoic glacial events and postglacial transgressions in Gondwana, in López-Gamundí, O.R., and Buatois, L.A., eds., Late Paleozoic Glacial Events and Postglacial Transgressions in Gondwana: Geological Society of America Special Paper 468, p. v–viii, doi: 10.1130/2010.2468(00). For permission to copy, contact [email protected]. ©2010 The Geological Society of America. All rights reserved.

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Glacial Episode III

Arabia Glacial Episode II

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A F R I C A

340 Ma Polar Path 360 Ma 340 Ma

INDIA

1 360 Ma

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AUSTRALIA 320 Ma

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Figure 1. Gondwana Supercontinent and the Late Paleozoic Ice Age (LPIA) basins with glacial evidence in their stratigraphic record highlighted. Glacial episodes after López-Gamundí (1997); polar path after Powell and Li (1994). Locations of contributions in this volume indicated by numbers corresponding to chapters in the volume.

we will be to understand the relationship between shifts in pCO2, temperature, and ice volume and greenhouse gas forcing of past and future climates. Additionally, a better understanding of the mechanisms of postglacial transgressions along basin margins will allow us to refine our search for fossil fuels through the identification of potential marine source rocks and coals. We hope this volume contributes to both ends. This volume’s contributions constitute a wide range of topics related to this extreme paleoclimatic episode in Earth’s history. Original presentations were part of the IGCP 471 project (“Evolution of the western Gondwana during the late Paleozoic: Tectonosedimentary record, paleoclimate and biological changes”)–sponsored session Late Paleozoic glacial events and postglacial transgressions in Gondwana during the 32nd International Geological Congress (Florence, August 2004). It was evident at that time that consensus had been reached on some basic problems about the LPIA, but some new challenges had emerged. Some of the unanswered questions that this volume attempts to address revolve around (1) relatively less known glacial deposits in some regions of Gondwana (Central Africa, Kreuser and Woldu, Chapter 4; Arabia, Melvin et al., Chapter 2); (2) the controversy between a single massive ice sheet versus numerous glacial centers and alpine glaciers (Isbell, Chapter 3);

(3) the chronostratigraphic resolution of paleofaunas (Cisterna and Sterren, Chapter 6) and paleofloras (Ianuzzi et al., Chapter 5) that coexisted with extreme glacial and relatively milder early postglacial conditions, and the presence of freshwater and brackish water ichnofaunas related to postglacial marine transgressions (Buatois et al., Chapter 7; Desjardins et al., Chapter 8); (4) the characterization of high-latitude, postglacial lakes (Miller and Isbell, Chapter 9); and (5) the search for a unifying sequencestratigraphic model for the glacial-postglacial transition (LópezGamundí, Chapter 1). The contributions included in this volume cover a broad geography across Gondwana, but they do not have the objective of giving a state-of-the-art review of the LPIA, a theme that has been periodically dealt with since Hambrey and Harland’s (1981) volume. A recent update can be found in Fielding et al. (2008). Rather, the present volume is focused on key specific topics related to the LPIA that, in our opinion, required further study. These topics deal with the two main episodes identified during the LPIA: the early Late Carboniferous (Namurian–Westphalian) glacial episode II, mostly confined to the Paleopacific margin of southern South America, and a much more widespread early Permian glacial episode III, which affected the rest of the supercontinent (López-Gamundí, 1997; Isbell et al., 2003). Instead of providing

Introduction brief summaries of specific areas, the authors were encouraged to expand their views, providing full documentation. The book consists of two main parts. The first half deals with sedimentologic, paleoenvironmental, and paleoclimatic aspects of the glacial event. The second half explores paleobiologic aspects of glacial and glacially influenced ecosystems. The first contribution, which is by López-Gamundí, focuses on the sequence stratigraphy of the late Paleozoic glacial event and the subsequent postglacial phase, setting the stage for the rest of the volume. He notes that, irrespective of the age, there is a common stratigraphic stacking pattern in each of the transgressive events. As a result of the combined effect of fast glacioeustatic sea-level rise and subsidence along basin margins, a drastic landward facies shift took place in the transition from glacially dominated to glacially influenced early postglacial environments. Available information allows recognition of two basic types of transgressive systems tracts (TSTs): (1) a complete TST, with a basal diamictite unit, followed by shale with ice-rafted debris (IRD) and IRD-free shales, culminating in a maximum flooding surface; and (2) a base-cut TST in which the TST is dominated by open-marine shales, generally devoid of IRD. The other three contributions are case studies based on specific areas of Gondwana but bearing implications at a more global scale. Melvin et al. (Chapter 2) provide a detailed characterization of the Upper Carboniferous–Lower Permian Unayzah Formation of subsurface Saudi Arabia. This unit (subdivided into four members) is particularly relevant, because it provides a full picture of the paleoclimatic evolution in this region of Gondwana, from glacial through postglacial to semiarid and arid conditions. The wide spectrum of facies documented includes tillite, reworked diamictite, glaciolacustrine fines and turbidites, fluvial deposits, paleosoils, and eolianites. These authors differentiate between climatically and tectonically controlled transgressions and provide correlations across the Arabian Peninsula. In Chapter 3, Isbell explores the paleoenvironmental and paleoclimatic implications of glaciomarine deposits in the Permian Metschel Tillite of the Transantarctic Mountains. In contrast to previous interpretations, he proposes that ice entered the area from at least two different ice centers on opposite sides of the basin. Abundant evidence of glaciotectonic structures is documented, including sheared diamictites and thrust sheets. The global significance of this study resides in that multiple glaciers contain less ice volume than a single massive ice sheet, impacting on global climate and eustatic sea level in a different way than would have a single massive ice sheet. Kreuser and Woldu (Chapter 4) provide a detailed characterization of Permian euxinic lake deposits preserved in the Idusi Formation of the Ruhuhu Basin in Tanzania. The succession reflects a transition from glacial to postglacial deposits, culminating in the development of distal alluvial fans during climatic amelioration. These extensive anaerobic stratified lakes provided appropriate conditions for deposition of black shales with abundant organic matter (up to 11% total organic carbon [TOC]). These authors underscore the importance of these deposits as source rocks and

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outline the regional extent of these anoxic lakes in various late Paleozoic basins of eastern and southern Africa. The nature of continental and shallow-marine ecosystems during glacial and postglacial times is still poorly understood. The last five contributions of this book focus on this topic, touching also on biostratigraphic implications. In Chapter 5, Iannuzzi et al. provide for the first time a sequence-stratigraphic and paleoenvironmental framework for palynozones and plant zones in the Rio Grande do Sul portion of the Paraná Basin. These authors find that the boundaries of palynozones lie near the maximum flooding surfaces. In addition, they note that the plant zones previously defined correspond to distinct ecofacies and are better regarded as ecozones rather than as biozones. Based on this analysis, a link is suggested between the increase in floral diversity of the Glossopteris-Rhodeopteridium Zone and the appearance of more complex coastal ecosystems as recorded in the Rio Bonito Formation. Cisterna and Sterren (Chapter 6) evaluate taphonomic and paleoecologic aspects of the Levipustula Fauna. This fauna is typical of lower Upper Carboniferous glacially related deposits in the Andean basins of Argentina. Based on studies in different stratigraphic units of the Calingasta-Uspallata Basin, these authors are able to distinguish two associations: intraglacial and postglacial. They also note that the postglacial association is more diverse and displays more abundance than the intraglacial association. This increase in diversity and abundance is explained as a result of less stressful conditions resulting from climatic amelioration. The last three contributions deal with ichnology. In particular, trace fossils are ideally suited for ecosystem studies because they provide in situ evidence of organism-substrate interactions. In Chapter 7, Buatois et al. summarize ichnologic data from eight different Gondwanan basins (Paganzo, San Rafael, Tarija, Paraná, Karoo, Falkland, Transantarctic, and Sydney) and note the presence of fresh-water ichnofaunas in direct association with glacially influenced coasts affected by strong discharges of meltwater as a recurrent theme. These authors suggest that freshwater conditions prevailed in coastal areas during most of the postglacial times because of a strong discharge of fresh water from melting of the ice masses. They conclude that the classic marine-nonmarine dichotomy used in ichnologic studies may be misleading in this type of setting. Desjardins et al. (Chapter 8) document the ichnofauna present in transgressive deposits of the uppermost Carboniferous and lowermost Permian Tupe Formation of the Paganzo Basin in western Argentina. These authors discuss ichnologic aspects of the transition from postglacial fluvial deposits to bay deposits formed during a rise in sea level. The recognized trace-fossil assemblages reflect the changing environmental conditions that result from a base-level rise. This study underscores the interplay of ecology, sea-level rise, and paleogeography as controlling factors for trace-fossil distribution. In Chapter 9, Miller and Isbell focus on the paleoecologic implications of an ichnofauna formed in a large and deep turbiditic Permian lake, recorded in the Mackellar Formation of the Transantarctic Mountains. The Mackellar ichnofauna is of low

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diversity, and the degree of bioturbation is generally low, suggesting oxic conditions and restriction of the benthos to areas with low rates of sedimentation. They compare this Permian lake with modern Lake Agassiz, and conclude that in spite of its high paleolatitude (~80°S), the lake was dynamic and biologically active. ACKNOWLEDGMENTS Finally the editors want to thank the following colleagues for dedicating their time to reviewing this volume’s contributions: Lucia Angiolini (Università degli Studi di Milano, Italy), Christoph Breitkreuz (Institut für Geologie, Freiberg, Germany), Roberto d’Avila (Petrobras, Brazil), Jim Collinson (USA), Almerio Barros França (Petrobras, Brazil), Dirk Knaust (StatoilHydro, Norway), Ricardo Melchor (Universidad de La Pampa, Argentina), Marcello Guimarães Simões (Sao Paulo State University at Botucatu, Brazil), Lynn Soreghan (University of Oklahoma, USA), Luis A. Spalletti (Universidad de La Plata, Argentina), Antonio Rocha-Campos (University of Sao Paulo, Brazil), Alfred Uchman (Jagiellonian University, Poland), John Veevers (MacQuarie University, Australia), and Joonas Virtasalo (University of Turku, Finland). REFERENCES CITED Crowell, J.C., 1978, Gondwana glaciation, cyclothems, continental positioning and climate change: American Journal of Science, v. 278, p. 1345–1372. Crowell, J.C., 1982, Continental glaciation through geologic time, in Climate in Earth History: Studies in Geophysics, Washington, D.C., National Academy Press, p. 77–82. Crowell, J.C., 1999, Pre-Mesozoic Ice Ages: Their Bearing on Understanding the Climate System: Geological Society of America Memoir 192, 106 p. Crowell, J.C., and Frakes, L.A., 1971a, Late Palaeozoic glaciation of Australia: Journal of the Geological Society of Australia, v. 17, p. 115–155. Crowell, J.C., and Frakes, L.A., 1971b, Late Paleozoic glaciation: Part IV, Australia: Geological Society of America Bulletin, v. 82, p. 2515–2540, doi: 10.1130/0016-7606(1971)82[2515:LPGPIA]2.0.CO;2. Crowell, J.C., and Frakes, L.A., 1972, Late Paleozoic glaciation: Part V, Karoo Basin, South Africa: Geological Society of America Bulletin, v. 83, p. 2887–2912, doi: 10.1130/0016-7606(1972)83[2887:LPGPVK ]2.0.CO;2. Du Toit, A.L., 1927, A Geological Comparison of South America with South Africa: Washington, D.C., Carnegie Institute, 157 p. Eyles, N., 1993, Earth’s glacial records and its tectonic setting: Earth-Science Reviews, v. 35, p. 1–248, doi: 10.1016/0012-8252(93)90002-O. Fielding, C.R., Frank, T.D., and Isbell, J.L., eds., 2008, Resolving the Late Paleozoic Ice Age in Time and Space: Geological Society of America Special Paper 441, 354 p.

Frakes, L.A., and Crowell, J.C., 1967, Facies and paleogeography of late Paleozoic Lafonian diamictite, Falkland Islands: Geological Society of America Bulletin, v. 78, p. 37–58. Frakes, L.A., and Crowell, J.C., 1969, Late Paleozoic glaciation: I, South America: Geological Society of America Bulletin, v. 80, p. 1007–1042, doi: 10 .1130/0016-7606(1969)80[1007:LPGISA]2.0.CO;2. Frakes, L.A., and Crowell, J.C., 1970, Late Paleozoic glaciation: II, Africa exclusive of the Karroo basin: Geological Society of America Bulletin, v. 81, p. 2261–2286, doi: 10.1130/0016-7606(1970)81[2261:LPGIAE ]2.0.CO;2. Frakes, L.A., Amos, A.J., and Crowell, J.C., 1969, Origin and stratigraphy of Late Paleozoic diamictites in Argentina and Bolivia, in Amos, A.J., ed., Gondwana Stratigraphy, IUGS Symposium (Buenos Aires, 1967): Earth Sciences, v. 2, p. 821–843. Frakes, L.A., Matthews, J.L., and Crowell, J.C., 1971, Late Paleozoic glaciation: Part III: Antarctica: Geological Society of America Bulletin, v. 82, p. 1581–1604, doi: 10.1130/0016-7606(1971)82[1581:LPGPIA ]2.0.CO;2. Hambrey, M., and Harland, W., 1981, Earth’s Pre-Pleistocene Glacial Record: Cambridge, UK, Cambridge University Press, 1044 p. Isbell, J.L., Miller, M.L., Wolfe, K.L., and Lenaker, P.A., 2003, Timing of late Paleozoic glaciation in Gondwana: Was glaciation responsible for the development of northern hemisphere cyclothems?, in Chan, M.A., and Archer, A.W., eds., Extreme Depositional Environments: Mega End Members in Geologic Time: Geological Society of America Special Paper 370, p. 5–24. López-Gamundí, O.R., 1997, Glacial-postglacial transition in the late Paleozoic basins of southern South America, in Martini, I.P., ed., Late Glacial and Postglacial Environmental Changes, Quaternary, Carboniferous–Permian and Proterozoic: Oxford, UK, Oxford University Press, p. 147–168. Montañez, I., Tabor, N.J., Niemeier, D., DiMichele, W.A., Frank, T.D., Fielding, C.R., Isbell, J.L., Birgenheier, L.P., and Rygel, M.C., 2007, CO2-forced climate and vegetation instability during late Paleozoic deglaciation: Science, v. 315, p. 87–91, doi: 10.1126/science.1134207. Powell, C.McA., and Li, Z.X., 1994, Reconstruction of the Panthalassan margin of Gondwanaland, in Veevers, J.J., and Powell, C.McA., eds., Permian– Triassic Pangean Basins and Foldbelts along the Panthalassan Margin of Gondwanaland: Geological Society of America Memoir 184, p. 5–9. Rygel, M.C., Fielding, C.R., Frank, T., and Birgeinheier, L.R., 2008, The magnitude of late Paleozoic glacioeustatic fluctuations: A synthesis: Journal of Sedimentary Research, v. 78, p. 500–511, doi: 10.2110/jsr.2008.058. Veevers, J.J., and Powell, C.M., 1987, Late Paleozoic glacial episodes in Gondwanaland reflected in transgressive-regressive depositional sequences in Euramerica: Geological Society of America Bulletin, v. 98, p. 475–487, doi: 10.1130/0016-7606(1987)98<475:LPGEIG>2.0.CO;2. Wanless, H.R., and Shepard, F.P., 1936, Sea level and climatic changes related to late Paleozoic cycles: Geological Society of America Bulletin, v. 47, p. 1177–1206. Wegener, A., 1915, Die Entsehung der Kontinente und Ozeane: Braunchweig, Germany, Vieweg, 367 p. Wright, V.P., and Vanstone, S.D., 2001, Onset of Late Paleozoic glacio-eustasy and the evolving climates of low latitude areas: A synthesis of current understanding: Journal of the Geological Society [London], v. 158, p. 579–582.

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The Geological Society of America Special Paper 468 2010

Transgressions related to the demise of the Late Paleozoic Ice Age: Their sequence stratigraphic context Oscar R. López-Gamundí Hess Corporation, 500 Dallas Street, Houston, Texas 77002, USA ABSTRACT The Gondwanan Icehouse Period spanned between the mid-Carboniferous and Early Permian waning by the early Late Permian. Early postglacial sea-level rise related to the final stage of the Late Paleozoic Ice Age favored creation of accommodation space with preservation potential for productive anoxia events in the newly inundated shelves and peat-forming conditions favored by rapid water table rise in updip positions in the basin. The combined effect of fast glacioeustatic sea-level rise and subsidence along basin margins led to a drastic landward facies shift; the newly created space was sufficient to accommodate a transgressive systems tract (TST) that, irrespective of the age of the glacial episode, exhibits common characteristics across Gondwana. High fresh-water discharges related to the retreat of glaciers resulted in associated reduction in coastal salinity. Therefore, fjord-like settings as part of early postglacial inland seas seem to be a valid analogue for many of these TSTs. The examples of glacial-postglacial transitions analyzed in this contribution are present in a variety of basin types, namely, those ranging from backarc foreland basins to rifts. In all of them a clear retrogradational stacking pattern is detectable in the transition from glacially dominated settings to glacially influenced early postglacial environments. Examples from South America (Calingasta-Uspallata and Paganzo Basins), South Africa (Karoo Basin), Peninsular India (several Gondwana basins), and eastern Australia (Tasmania Basin) help define two basic types of TSTs: (1) complete TSTs, with a basal part of clast-poor, massive to poorly stratified diamictites, thinly bedded diamictites, shales with ice-rafted debris (IRD) and IRD-free shales, and an upper part dominated by open-marine shales representing the maximum flooding of the shelf; and (2) base-cut TSTs in which the basal transgressive portion is mostly omitted, and the TST is exclusively represented by open-marine shales generally devoid of IRD. Whereas the complete TSTs are common in cases in which high sediment supply rates via rain-out, ice rafting, and settling of fines prevail during the early phase of deglaciation, the base-cut TSTs, on the other end, reflect the dominance of drastic sealevel rises related to fast glacier retreats.

INTRODUCTION

the Earth (Fig. 1). Evidence of this glaciation is widespread in the Gondwana Supercontinent (Hambrey and Harland, 1981; Crowell, 1983, 1999; Eyles, 1993; Fielding et al., 2008a). This glacial age coincides with a high paleolatitude for Gondwana and the

The Late Paleozoic Ice Age (LPIA) is one of the best recorded episodes of extreme climatic conditions that affected

López-Gamundí, O.R., 2010, Transgressions related to the demise of the Late Paleozoic Ice Age: Their sequence stratigraphic context, in López-Gamundí, O.R., and Buatois, L.A., eds., Late Paleozoic Glacial Events and Postglacial Transgressions in Gondwana: Geological Society of America Special Paper 468, p. 1–35, doi: 10.1130/2010.2468(01). For permission to copy, contact [email protected]. ©2010 The Geological Society of America. All rights reserved.

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Figure 1. Reconstruction of Gondwana supercontinent with simplified outlines of principal basins during the Late Carboniferous–Early Permian (glacial episodes II and III). Positions of Arabia and Madagascar after De Wit et al. (1988); South American basins after López-Gamundí et al. (1994); African basins from Veevers et al. (1994) and Visser (1997a); Falkand-Malvinas Islands in their pre-breakup position east of the coast of South Africa as originally proposed by Adie (1952); Indian basins after Wopfner and Casshyap (1997); Australian basins modified from Struckmeyer and Totterdell (1992) and Lindsay (1997). 1—Calingasta-Uspallata and Paganzo; 2—Tarija; 3—San Rafael; 4—Tepuel-Genoa; 5—Sauce Grande; 6—Chaco-Paraná; 7—Paraná; 8—Huab; 9—Kalahari; 10—Karoo; 11—Falkland-Malvinas Islands; 12—Zambesi; 13—Congo; 14—Tanzania; 15—Malagasy; 16—Peninsular India; 17—Extra-Peninsular India (Himalayan); 18—Salt Range; 19—Yemen–South Arabia; 20—Oman; 21—Pensacola Mountains; 22—Transantarctic Mountains; 23—Tasmania; 24—Murray; 25—Sydney; 26—Bowen; 27—Galilee; 28—Cooper; 29—Pedirka-Arckaringa; 30—Perth; 31—Carnarvon; 32—Canning; 33—Browse; 34—Bonaparte.

growth of high standing topography when Gondwana collided with Laurasia to create Pangea (Eyles, 2008). The stratigraphic record derived from this glaciation is abundant and diverse: different types of glacioterrestrial and glaciomarine sediments associated with glacially abraded surfaces are present in a wide variety of basin types and tectonic settings. Despite this variability of facies, basin types, and tectonic regimes, most of these glaciated margins were affected during the early deglaciation phase by a sea-level rise that imparted a series of conspicuously diagnostic stacking patterns that can be interpreted in sequence stratigraphic terms. Examples of such glacial-postglacial transition have been documented not only for the LPIA but also for the Pleistocene glaciation. Owing to its relation with large hydrocarbon reservoirs, evidence of such transition from surface and subsurface studies has been abundantly documented for the Late Ordovician glaciation that affected most of the North African and Arabian platforms (Beuf et al., 1971; Deynoux et al., 1985; Vaslet, 1990;

Ghienne et al., 2007). The objective of this contribution is to illustrate the sequence stratigraphic context of the transgressions in relation to the demise of the LPIA. The emphasis of this study is on the section temporally constrained to the initial deglaciation when glacioisostatic rebound lags behind a rapid eustatic rise in sea level. CHRONOSTRATIGRAPHIC FRAMEWORK The record of the LPIA seems to cluster around three distinct episodes (Veevers and Powell, 1987; López-Gamundí, 1997; Isbell et al., 2003). The Late Devonian is the oldest of these episodes (episode I of López-Gamundí, 1997); it was originally described by Caputo (1985) and is apparently confined to a broad SW-NE belt across the central and northern parts of South America (Caputo el al., 2008; López-Gamundí and Buatois, this volume) from Peru (Carlotto et al., 2004; Cerpa et al., 2004) and

Transgressions related to the demise of the Late Paleozoic Ice Age Bolivia (Díaz Martínez and Isaacson, 1994) to Brazil (Solimões Basin, Eiras et al., 1994; Amazonas basin, Cunha et al., 1994; Parnaiba, Goes and Feijo, 1994) and recently extended to the Central African Republic and Niger in Africa (Isaacson et al., 2008). A late Fammenian age has been suggested for this glacial event on the basis of palynological evidence mostly from Bolivia (Isaacson et al., 1999). A northward extension of this short-lived, areally confined glaciation has been proposed for North America (Cecil et al., 2004; Isaacson and Díaz Martínez, 2005; Brezinski et al., 2008), but its evidence is tenuous so far and requires further work. The postglacial transgressive sections analyzed herein correspond to glacial episodes II and III as defined by López-Gamundí (1997). The former is associated with the mid-Carboniferous Levispustula fauna, and the latter is characterized by the Early Permian (Sakmarian) Eurydesma fauna. These two cold-water faunas have been identified in the marine fine-grained sediments that rest directly on, or in similar facies interbedded with, glaciogenic deposits across the Gondwana Supercontinent. The age of the Levipustula fauna has been traditionally considered Namurian–Westphalian (Roberts et al., 1976; see Cisterna and Sterren, this volume, for a recent review on the age of this fauna) but more recently has been confined to the early Namurian on the basis of absolute ages from SHRIMP (sensitive high-resolution ion microprobe) zircon-based dating in interbedded tuffs in Australia (Roberts et al., 1995; Fielding et al., 2008b). Glacial episode II is restricted to the westernmost part of the Gondwanan Supercontinent in southern South America (Amos and López-Gamundí, 1981a; González, 1990; López-Gamundí, 1984, 1989, 1997). Glacial episode III is widespread across the Gondwanan basins from South America (Frakes and Crowell, 1969; Rocha-Campos and dos Santos, 1981; López-Gamundí, 1997; Rocha-Campos et al., 2008) across Africa (Crowell and Frakes, 1972; Theron and Blignault, 1975; Visser, 1983, 1987a, 1987b, 1989, 1997a, 1997b; Wopfner and Kreuser, 1986; Von Brunn, 1994, 1996), Arabia (Helal, 1964; Kruck and Thiele, 1983; McClure, 1980; Braakman et al., 1982; Levell et al., 1988; Melvin and Sprague, 2006; Melvin et al., this volume) to India (Niyogi, 1961; Smith, 1963; Casshyap and Qidwai, 1974; Ghosh and Mitra, 1975; Casshyap and Srivastava, 1987), and Antarctica (Lindsay, 1970; Miller, 1989; Collinson et al., 1994; Isbell et al., 2008) to Australia (Crowell and Frakes, 1971; Struckmeyer and Totterdell, 1992; Lindsay, 1997). SHRIMP ages from the lower Dwyka Formation (Bangert et al., 1999) in the Karoo Basin (Fig. 1) and the Itararé Group (Rocha-Campos, 2006, in Rocha-Campos et al., 2008) in the Paraná Basin (Fig. 1) indicate a Late Carboniferous (Stephanian) age for the onset of glacial episode III in South Africa and South America. There seems to be consensus on a Sakmarian age for the widespread Early Permian collapse of ice sheets (Dickins, 1996; Isbell et al., 2003). Tuff zones at the base of the postglacial Prince Albert Formation in the southwest Karoo Basin provided SHRIMP ages of ca. 290 Ma (Bangert et al., 1999), a mid-Sakmarian age based on the numerical time scale of Gradstein et al. (2004). Exceptionally, local glacial conditions per-

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sisted after the waning of the Gondwanan Ice Sheet Complex (Eyles, 1993) until the end of the Early Permian (Artinskian) along some margins and upland regions of western Gondwanan basins (northern margins of the Karoo Basin and the Kalahari Basin, Fig. 1) and particularly in eastern Gondwana, where small alpine ice caps provided glacial debris during three more minor glacial episodes in eastern Australia (Jones and Fielding, 2004). The proof of these waning glacial conditions elsewhere in Australia is mostly derived from the presence of ice-rafted debris (IRD) in the form of isolated dropstones in the Tasmania Basin (Fig. 1, Clarke and Banks, 1975) and the Sydney Basin (Fig. 1, Eyles et al., 1998). Indirect evidence for the presence of these waning, post-Sakmarian cold periods in eastern Australia is provided by studies that show lowered atmospheric pCO2 before the permanent transition to an ice-free Earth ca. 260 Ma (Montañez et al., 2007). The glacial-postglacial transition exhibits a similar retrogradational stacking of facies irrespective of their ages; therefore the analysis attempted in this contribution will be focused on the sedimentological characteristics of these transgressive deposits and their sequence stratigraphic context rather than on their chronostratigraphic significance. Analogies for vertical stacking patterns, for example, will be highlighted irrespective of the ages of the successions. Furthermore, probably the best documented pre-Pleistocene glacial-postglacial transition is the one that corresponds to the Late Ordovician (Hirnantian, late Ashgill) glaciation in North Africa and the Arabian Platform (Beuf et al., 1971; McClure, 1978; Hambrey, 1985; Vaslet, 1990; Le Heron et al., 2007; Ghienne et al., 2007), which shares sedimentological and sequence stratigraphic similarities with the late Paleozoic examples analyzed herein. Both glacial ages share a similar pattern of deglaciation characterized by a drastic landward facies shift during a fast transgression. However, unlike the short-lived Late Ordovician glaciation restricted between ~1 m.y. (Brenchley et al., 1994) and 10 m.y. (Ghienne, 2003), the LPIA is the longest lived glaciation in Earth history, perhaps one to two orders of magnitude longer than the Late Ordovician event (Eyles, 1993). Thus, eustatic fluctuations caused by numerous ice advances and retreats are commonly recorded across Gondwana for the late Paleozoic. The emphasis in this contribution is on the sedimentological and sequence stratigraphic characteristics of the wellpreserved glacial-postglacial transition rather than on the shorter lived, often cannibalized, glacial-interglacial fluctuations. Four examples have been selected; they cover a wide spatial and temporal range from west to east: Calingasta-Uspallata and Paganzo Basins (glacial episode II), Karoo and associated basins (glacial epsiode III), basins in Peninsular India (glacial episode III), and the Tasmania Basin in Australia (glacial episode III) (Fig. 1). SEQUENCE STRATIGRAPHIC CONTEXT FOR POSTGLACIAL TRANSGRESSIONS The term glacioeustasy is referred in this contribution to the process that generates changes in sea level that can be related to

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changes in ice volume. A first approximation to the magnitude of eustatic sea-level rises related to the rather abrupt demise of any glaciation is provided by estimates of volumes of ice released. Table 1 is a compilation of the estimated present-day area and volume of glaciers and the maximum sea-level rise potential. It is worth noting that minor melting episodes related to small, local ice caps have negligible effects on sea-level fluctuations, as opposed to major waxing of large areas (i.e., East Antarctica). In the latter scenario, significant (~65 m) maximum sea-level rises are expected. This effect could have been attenuated if several small ice sheets, rather than a large single one, were present as suggested for at least part of the LPIA by Isbell et al. (2003). A similar approach has been used by Le Heron and Dowdeswell (2009) for their calculation of the postglacial sea-level rise related to the Late Ordovician glaciation. These authors estimated a postglacial eustatic sea-level rise of ~75 m on the basis of a small ice-sheet hypothesis and concluded that this approach adequately accounts for the estimated magnitude of postglacial transgression (~45–80 m) associated with ice-mass decay. Their estimate of 50 m of postglacial sea-level rise only from the contribution of an ice sheet in the North African Platform (the largest of the three ice sheets modeled) is in agreement with independent estimates derived from studies on depth-related benthic fossil communities (Ross and Ross, 1996). The effects of isostasy and its lateral variations along a continental margin (i.e., differential cross-shelf isostatic response of a thin crust) may potentially mask the otherwise dominant effect of glacioeustatically induced sea-level rises on the stratal architecture and sequence development of the early postglacial basin fill. This was frequent in tectonically active regions affected by glaciation, in which the individual factors related to the glaciation (eustasy and isostasy) and those of local origin (i.e., tectonically induced subsidence) are difficult to discern. Studies of Pleistocene glaciation along the western Canadian continental shelf show this extreme variability (Clague et al., 1982). Whereas a marine transgression owing primarily to eustatic rise occurred in the southern regions of the Strait of Georgia adjacent to the central and northern Strait of Georgia, it appears that 100 m of isostatic adjustment associated with local tectonic changes were

offset by eustatic effects resulting in minimal sea-level fluctuations (Barrie and Conway, 2002). Nevertheless, the Pleistocene and pre-Pleistocene glacial-postglacial stratigraphic record provides abundant examples of deepening-upward sections that indicate a drastic increase in accommodation space during the early stages of deglaciation. This record is confined to a portion of the basin where the postglacial sea-level rise and the isostatic subsidence caused by water loading might have exceeded the combined effect of glacial erosion and glacio-isostatic rebound (Bjorlykke, 1985; Nystuen, 1985), resulting in a subsequent glacioeustatic transgression (Crowell, 1978). Areas close to the ice margin are commonly dominated by isostatic depression and rebound, whereas more distal areas are dominated by eustatic changes like submergence during deglaciation (Miller, 1996). The main controls on stratal architecture in any given sequence can be narrowed down to (1) accommodation space (the summation of subsidence rates and sea-level fluctuations), and (2) sediment supply. Figure 2 illustrates some of the possible conditions under which accommodation space is created during a stage of sea-level rise in a subsiding basin (Jervey, 1988). Case A illustrates a classic example in sequence stratigraphy where the summation of a constant (in this case slow to moderate) subsidence rate and a symmetrical sinusoidal eustatic sea-level curve creates a relative sea-level curve (accommodation space). Cases B, C, and D illustrate three alternative outcomes of the most likely scenario for postglacial transgressions when the eustatic sea-level cycle shows a clear asymmetry from a faster (glacioeustatic) sea-level rise (expressed by the steep part of the eustatic sea-level curve). The subsidence rates vary and the eustatic sealevel cycle remains constant in these three last cases (Figs. 2B, 2C, and 2D). Case C differs from case B only in that case C includes a faster subsidence rate; in this situation the postglacial sea-level rise could be partially masked because very fast subsidence rates exceed the sea-level rate with accommodation space not significantly destroyed, and consequently, even when eustatic fall is taking place, deep-water conditions similar to those dominant during the transgressive stage may persist in the highsea-level stage. Finally, case D illustrates an extreme case with variable subsidence rates (from initially moderate to fast) and

TABLE 1. ESTIMATED PRESENT-DAY AREA AND VOLUME OF GLACIERS AND MAXIMUM SEA-LEVEL-RISE POTENTIAL Percent Volume Percent Maximum sea-level-rise Geographic region Area 2 3 (%) (km ) (%) potential (km ) (m) Ice caps, ice fields, valley glaciers, etc. 680,000 4.24 180,000 0.55 0.45 Greenland (inland ice) 1,736,095 10 .82 2,600,000 7 . 90 6. 5 0 Local ice caps and other glaciers 48,599 0.30 20,000 0.06 0.05 Antarctica 13,586,400 84.64 30,109,800 91.49 73.44 East Antarctica 10,153,170 26,039,200 64.80 West Antarctica 1 ,9 1 8 , 1 7 0 3,262,000 8.06 Antarctic Peninsula 4 46,690 227, 100 0. 46 Ross Ice Shelf 53 6 , 0 70 2 2 9, 6 0 0 0 .01 Ronne-Filchner ice shelves 532 ,20 0 351 ,90 0 0.11 Totals 16 , 0 5 1 , 0 9 4 100 .0 0 32,909,800 1 00 .0 0 8 0. 4 4 Notes: From Williams and Ferrigno (2008). Values in italicized rows are a subset of the “Antarctica” row values.

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creation of deep-water conditions during transgression and even deeper water conditions during the high-sea-level stage. In cases C, and particularly D, subsidence rates are so high that the basin undergoes no decrease in accommodation even though eustatic fall may be occurring (Emery and Myers, 1996). Cases C and D correspond to the concept of forced transgression of Chough and Hwang (1997). The classic definition of sequence is adopted in this contribution. Sequences are composed of depositional packages or systems tracts (Brown and Fisher, 1977) deposited during specific phases of the relative sea-level cycle (Posamentier and Allen, 1999). A systems tract is a linkage of contemporaneous depositional systems and is defined by the nature of its boundaries (see below). Three principal tracts are used in this contribution: lowstand, transgressive, and highstand systems tracts (LST, TST, and HST, respectively). Boundaries between system tracts are defined by key sequence-stratigraphic surfaces: 1. Sequence boundary (SB): an unconformity and its correlative conformity corresponding to the most regressive configuration of stratigraphic architecture. A sequence boundary is expressed as a facies dislocation, a surface where rocks of a shallower facies rest directly on rocks of a significantly deeper facies (Emery and Myers, 1996). This facies dislocation (basinward facies shift) implies a fall in relative sea level. 2. Flooding surface (FS): a surface that separates abruptly deeper deposits from underlying shallower water sediments (Van Wagoner et al., 1990). This surface implies a rise in relative sea level and can be generated by eustasy,

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Figure 2. Accommodation space as a function of subsidence rates and sealevel changes. Relative sea level (RSL) is equivalent in this case to accommodation because both curves begin at zero-depth water. Modified from Jervey (1988) and Emery and Myers (1996). (A) Accommodation space curve as the summation of a sinusoidal eustatic sealevel curve (ESL) and a slow to moderate subsidence rate. B, C, D: Resulting accommodation space from an asymmetric sinusoidal sea-level curve (ESL) and variable subsidence rates. (B) Slow to moderate subsidence (same as in A). (C) Moderate to high subsidence rate; note higher resulting accommodation space than in B; note that with faster subsidence, maximum accommodation is progressively later. (D) Increasing subsidence rate, which creates the most accommodation space. See discussion in text.

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tectonics, and/or delta-lobe switching; when recognizable globally, a flooding surface can be assigned to a eustatic sea-level rise. 3. Maximum flooding surface (MFS): the surface that corresponds to the most transgressive stratigraphic architecture. Maximum flooding surfaces can be identified by the presence of condensed sections that reflect distant sediment sources at the peak of transgression (Loutit et al., 1988; Posamentier and Allen, 1999). In shelfal and basinal settings, maximum flooding surfaces can be identified by evidence of condensation (i.e., firmgrounds), fine-grained (silt-clay) deposits, the presence of high organic content (expressed as high gamma-ray values in well logs), phosphate levels, fossiliferous beds, and outer shelf carbonates. In paralic successions the MFS is coeval with the most landward position of the shoreline (Emery and Myers, 1996). In areas where it is difficult to identify and trace laterally a maximum flooding interval (MFI) or zone is recognized for such a surface. The MFS should lie within the MFI, but no sedimentological evidence could be found to define such a surface. 4. Transgressive ravinement surface (TRS): the erosional surface cut by wave action during transgression. This transgressive surface is the first significant marine flooding surface across the shelf within a sequence (Van Wagoner et al., 1988). It defines the base of the TST. As pointed out by Posamentier and Allen (1999), some potential confusion arises when terms like transgressive surface and flooding surface are used because they are commonly

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López-Gamundí used interchangeably. The former is reserved in this contribution for a surface “marking the onset of significant and extended period of transgression within a succession” (Posamentier and Allen, 1999, p. 5), restricting flooding only to the process of subsequent inundation.

FACIES ASSOCIATIONS RELATED TO POSTGLACIAL TRANSGRESSIONS With few exceptions in which terrestrial subglacial tills and associated facies are the dominant deposits, the stratigraphic record of the LPIA and its subsequent final deglaciation record across Gondwana have been preserved mostly in subaqueous settings. The main reason for the skewed record is possibly the significantly different preservation potential of terrestrial and subaqueous (marine or lacustrine) glacial environments. Preservation potential depends on a delicate balance of long- and shortterm components. The long-term component is largely controlled by the geotectonic setting (i.e., forearc basin, cratonic basin, etc.) and the related tectonic activity of the basin after deposition (i.e., postdepositional erosion); the short-term component is connected to the erosion coeval with sedimentation (Nystuen, 1985). This short-term component of the preservation potential is significantly low for terrestrial settings. Two main facies associations were connected with glacier retreat by López-Gamundí (1997) based on either nonmarine or marine conditions that prevailed in any specific part of a basin. Those are, respectively, a valley-glacier-retreat facies association and a submarine-retreat facies association. The Eyles et al. (1983) facies code for diamictite sequences is used in this contribution. Valley-Glacier-Retreat Facies Association This facies association developed along basin margins under predominantly subaerial or shallow subaqueous conditions during the retreat of a glacier front. Owing to the relatively low initial subsidence along the basin margins, the fill of these valleys is modest in thickness when compared with the submarine-retreat facies association developed basinward. Any sea-level rise has significant effects inland by raising the water table. This influence of a sea-level rise on mires was summarized by Bohacs and Sutter (1997), who estimated that a 5 m rise in sea level would propagate inland 50 km from the shore across a low-topography coastal area, producing a 3.5 m rise in the groundwater table. Terrigenous organic matter can be preserved to form coals only when and where the overall increase in accommodation approximately equals the production rate of peat (Bohacs and Sutter, 1997). For mires, base level is the groundwater table. Subsidence varies much more slowly than does the groundwater table and thus is commonly the major long-term control of accommodation in nonmarine settings. The postglacial facies in the valley-glacier-retreat setting rests on basement or, more commonly, on glaciogenic facies represented by thin (5–10 m), laterally discontinuous, massive

to crudely stratified diamictite (interpreted as subglacial till) beds partly associated with lacustrine shales with IRD. The glacial beds grade upward to fluvial deposits, coals, and carbonaceous shales, which represent the postglacial stage. The typical fluvial facies are predominantly coarse-grained sediments made up of pebble to cobble conglomerates, pebbly sandstones with trough cross-bedding, and medium- to fine-grained rippled sandstones deposited in a gravelly braided channel complex or as part of a subaqueous proximal outwash fan during a phase of glacier retreat. In coastal environments, marine influence can be expressed as marine fossiliferous shales, considered as maximum flooding surfaces, whereas the underlying coals and carbonaceous rocks are interpreted as initial flooding surfaces within a TST. In exclusively nonmarine settings the coals may represent the correlatives of maximum flooding surfaces developed basinward. Coals and carbonaceous shales present in this facies association have several characteristics that can be related to the accumulation of peat on initial rapid transgressions and increasing accommodation space. They are relatively thin and areally restricted as a result of increasing accommodation (cf. Bohacs and Sutter, 1997). Also, they are characterized by high sulfur (S) content since the original mires near the coast may have been submerged in brackish or salt water. In peat forming environments with brackish and marine influence (i.e., rheothropic mires), groundwater and seawater are important sources of primary S. In these brackish- to marine-influenced peats, the S content can be high (>10% in extreme cases), whereas in fresh-water peats tend to be low (<1%) (Casagrande et al., 1977; Cohen et al., 1989; Phillips and Bustin, 1996). Therefore, high S contents in coals have been traditionally used to infer marine influence (Williams and Keith, 1963; Raymond and Davies, 1979). Submarine-Retreat Facies Association This facies association is present in basin margins under waning glacial marine influence; it is characterized by the dominance of fine-grained (clay-silt) sedimentation in open-marine areas predominantly below wave base. Underlying glacial facies are characterized by laterally extensive beds of massive to poorly stratified diamictites (Dmm and Dms, Eyles et al.,1985) interpreted as rain-out tills, thick to thin-bedded diamictites interpreted as the product of remobilization of glacial debris via gravity flows (mostly discrete debris flows), and IRD-bearing mudstones and thin-bedded diamictites with outsize clasts (dropstones), interpreted as the combined product of cohesive debris flows and ice rafting (López-Gamundí, 1991). The presence of intra-till striated boulder pavements within mostly massive diamictites (interpreted as subglacial tills) is common and reflects high-frequency advance-retreat fluctuations of the glacier (cf. Eyles, 1988). These underlying glacial facies are similar to those described by Evans and Pudsey (2002) as proximal glaciomarine sediments under the Antarctic Peninsula ice shelf. They consist of coarse-grained facies with stratification and laminae, gradational

Transgressions related to the demise of the Late Paleozoic Ice Age contacts, dropstone structures (cf. Ovenshine, 1970; Thomas and Connell, 1985) coarse-tail grading, and/or sorted sand and silt. Facies associations indicate the interaction of three sedimentary processes following decoupling of grounded ice during deglaciation: meltout, gravity flow, and current transport. This proximal zone is an area of high sedimentation owing to ice melting (Drewry and Cooper, 1981; Powell, 1984); therefore, rain-out tills prevail (Powell et al., 1986). Deposits from gravity flows include bedded diamicton, internally massive with normal coarse-tail grading and with sharp contacts. These characteristics indicate deposition from cohesive debris flows. Similar thin-bedded (<50 cm) debris flow deposits have been described in Pleistocene glacially influenced environments such as those around the continental margin of Antarctica (Wright et al., 1983), the Nova Scotian Slope (Hill et al., 1982), and Baffin Bay (Aksu, 1984), and they constitute a conspicuous facies in the early postglacial TSTs. The fine clastics of this facies association can rest directly over the glaciomarine deposits or grade upward from an intervening facies characterized by mudstones with IRD and thinbedded diamictites. This transitional facies is interpreted as the final stage of sedimentation under glacial influence as evidenced by glacially derived gravity (debris) flow deposits and ice rafting. The glacial-postglacial section exhibits a clear fining-upward tendency in concert with a drastic facies shift typical of conditions when the rate of accommodation (in this case dominated by the rate of sea-level rise) exceeds the sediment supply rate (submarine retreat model, Edwards, 1986). A modern analogue of this depositional setting is present in the ice shelves of the Antarctic Peninsula (cf. Evans and Pudsey, 2002), where massive, bioturbated terrigenous silty clay dominates sedimentation as the grounding line recedes toward the coast with progressive deglaciation. Decreasing quantities of IRD may also indicate significant distance from the retreating glacier front. EXAMPLES OF POSTGLACIAL TSTS IN GONDWANA Calingasta-Uspallata and Paganzo Basins The Calingasta-Uspallata and Paganzo Basins lie along the western margin of Gondwana (Fig. 1). They are part of a series of backarc foreland basins developed along the active Paleopacific margin (López-Gamundí et al., 1994). A series of basement highs fragmented sedimentation in this basin with a depocenter close to the arc on the west (Calingasta-Uspallata) and a segmented basin cratonward (Paganzo) (Fig. 3). Sedimentation began in about the Early Carboniferous (Sessarego and Cesari, 1989) and expanded during the mid-Carboniferous (Namurian) with predominantly glacially influenced marine sedimentation in the Calingasta-Uspallata Basin and the western Paganzo Basin (Frakes and Crowell, 1969) as well as transitional environments in the eastern part of the Paganzo Basin. Glacial deposits are widespread in both basins; they correspond to glacial episode II (López-Gamundí, 1997), which is biostratigraphically confined by the presence of the Levipustula fauna mostly in overlying

7

postglacial shales in the Calingasta-Uspallata Basin (Amos and Rolleri, 1965; González, 1985). An increasing body of evidence suggests that sedimentation in the western and parts of the eastern Paganzo Basin took place in an inland sea at least during the transgression that followed the glacier retreat, creating conditions for fjord-like sedimentation with variable salinity (Kneller et al., 2004; Buatois et al., 2006; Desjardins et al., this volume; Limarino et al., 2010). Glacial sedimentation in the Calingasta-Uspallata Basin and the western Paganzo Basin was dominated by massive (Dmm) to poorly stratified (Dms) sandy or muddy diamictites, interpreted as subglacial and rain-out tills, resting on metasedimentary basement, associated with well-stratified muddy diamictites (Dms[r]), pebbly (dropstone) shales (Fld), and high- and lowdensity turbidites (López-Gamundí, 1987). Evidence of glacial abrasion in both basins is widespread (López-Gamundí and Martínez, 2000). The basal fill of the Calingasta-Uspallata Basin and the western Paganzo Basin exhibits significant lateral facies changes caused by progressive sedimentary fill across an irregular pre-Carboniferous topography. In the Calingasta-Uspallata Basin the basal deposits range from proximal glaciomarine deposits with intra- and inter-till striated boulder pavements to glacially influenced sediments deposited by sediment gravity flow, mostly cohesive debris flows, with high- and low-density turbidity currents. These facies are widespread along the eastern margin of the Calingasta-Uspallata Basin (Fig. 3). The Hoyada Verde section on the eastern basin margin (Fig. 4) is an example of the glacial-postglacial transition. The lower glacial section is composed of clast-rich, massive to crudely stratified diamictites (Dmm) interpreted as subglacial tills (Fig. 5A), stratified, clast-poor diamictites (Dms) considered rain-out tills, and subordinate thinly bedded diamictites with IRD, interpreted as the combined product of debris flows and ice rafting. This glacial section is capped by a single-layer boulder pavement (Fig. 5B) made up of clasts up to boulder size with parallel to subparallel striations (González, 1981; López-Gamundí, 1983). The boulder pavement caps a thick, mostly massive diamictite section with isolated paleochannels containing reworked fossils and glendonite nodules, as reported by González (1981), who also identified glendonite concretions in underlying shales within the glaciogenic section. Paleo–ice flow directions are from NNE to SSW, coincident with paleocurrents obtained from ripples in the postglacial fine-grained sandstones (Fig. 4, López-Gamundí, 1983). The pavement is overlain by a laterally discontinuous, thin (>20 cm) bed of massive muddy diamictite (López-Gamundí and Martínez, 2000) that passes upward into thin-bedded diamictites (pebbly mudstones) with outsize clasts (Fig. 5C) and shales with dropstones that are considered part of the early postglacial TST. This facies association grades upward by gradual disappearance of the thin-bedded diamictites with IRD into an interval dominated by IRD-free shales (Fig. 5D) with marine fossils of the Levipustula fauna (Sterren and Cisterna, 2006; Cisterna and Sterren, this volume).These fossiliferous, IRD-free shales are considered the maximum flooding interval (MFI) in the late postglacial

Si S

G

Si

SG

SEDIMENTARY DOMAINS

Open marine Transitional (inland sea) Continental

Si

SG

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H

C

Pa

M

C Si

HV

SG

Pa

Je

Si

SG

AJ 100 km Coarse-grained sandstones Fine-grained sandstones with ripples (Sr) Fine-grained sandstones Mudstones

TST Late

Shales (Fl)

Early

Dropstone shales (Fld) Thinly bedded diamictites (Dms) HV

AJ

Massive and crudely stratified diamictites (Dmm)

20 m

0

Figure 3. Calingasta-Uspallata and Paganzo Basins during glacial episode II (López-Gamundí, 1997) with sections illustrating the glacial-postglacial transition in the three main sedimentary domains. Agua de Jagüel (AJ) section modified from López-Gamundí (1984); Hoyada Verde (HV) section modified from López-Gamundí (1983); Malanzán (M) section from Andreis et al. (1986); Cortaderas (C), Huaco (H), and Paganzo (Pa) sections after Limarino et al. (2002). Other location cited in the text: Jejenes (Je). Only upper-most glacial section illustrated in logs. Grain-size scale in logs: Si—silt; S—sand; G—gravel. TST— transgressive systems tract.

TS

HST

Figure 5D

Maximum Flooding Interval (MFI)

Early TST

H o y a d a Ver d e F o r m a t i o n

Figure 5C

TST

Figure 5B

Figure 5A

30 m

0 Si S

Shales (Fl)

Trough crossstratification

Mudstones

Starved ripples

Dropstone shales (Fld)

Burrows

Fine-grained sandstones

Marine fauna (Levipustula)

Pebbly sandstone

Deformed sandstone bodies

Conglomerates

Dropstones (IRD)

Stratified diamictites (Dms), thinly bedded

Paleo ice flow directions

Stratified diamictites (Dms)

Striated boulder pavement (SBP)

Massive diamictites (Dmm)

Paleocurrents

G

Figure 4. Hoyada Verde Formation section in Calingasta-Uspallata Basin. Adapted from López-Gamundí (1983, 1984). Grain-size scale in log: Si—silt; S—sand; G—gravel. HST—highland systems tract; TS—overlying Tres Saltos Formation; TST—transgressive systems tract.

10

López-Gamundí

A

C

B

D Sr+Fl

Fl

Figure 5. Hoyada Verde Formation outcrops; see log in Figure 4 for stratigraphic locations of photos. (A) Clast-rich massive diamictite (Dmm); circled hammer as scale in center. (B) Inter-till striated boulder pavement. (C) Thin-bedded diamictites (Dms) with dropstone, early TST. (D) Postglacial fossiliferous shales (Fl) of the late TST and overlying fine-grained sandstones and mudstones (Sr+Fl), part of the HST; cliffforming sandstone beds on the left (west) correspond to the overlying Tres Saltos Formation.

TST. The upward transition to the HST (Fig. 4) is manifested by increasing participation of interbedded mudstones and finegrained sandstones with slightly asymmetric to symmetric (wave) ripples (Figs. 4 and 5D; “flagstones” of Mésigos, 1953) with a clear coarsening- and thickening-upward stacking pattern. Traces common in marginal-marine settings were recognized by Mángano et al. (2003). This interval has been interpreted as deposited in the transition from offshore to lower shoreface (Cisterna and Sterren, this volume); it is considered in this contribution as part of the HST (Fig. 4). The predominant glacial association in the contiguous Paganzo Basin consists of massive and crudely bedded diamictites (Dmm), thin-bedded diamictites (Dms), and shales with ice-rafted material (dropstone facies, Fld) deposited in a brackish seaway (Guandacol embayment) partially open to the CalingastaUspallata Basin (Fig. 3). The western part of the Paganzo Basin was more influenced by marine conditions, and several examples of previously interpreted lacustrine mudstones resting on glacial deposits are now considered part of the postglacial transgression in coastal fjord-like environments (Limarino et al., 2002;

Kneller et al., 2004; Limarino et al., 2004, 2010) after a marine microfauna was reported from various localities (i.e., Las Lajas Formation, Je in Fig. 3, Cesari and Bercowski, 1997; Guandacol Formation, Ottone, 1991). The microfauna found in the postglacial deposits of the Guandacol Formation is associated with low-salinity linguliformean brachiopods (Martínez, 1993) and a brackish water ichnofauna (Buatois et al., 2006; Desjardins et al., this volume). Two main ichnofacies separated by a nonbioturbated black shale were identified. A lower depauperate Cruziana assemblage is characterized by brackish-water, low-diversity, simple forms in fine-grained sandstones with current and combined-flow ripples and IRD, and a granule conglomerate and very coarse grained sandstones (Mángano and Buatois, 2004). The upper assemblage (fresh-water Mermia ichnofacies, Buatois and Mangano, 2003) is present above the nonbioturbated black shales (MFI) and is comparatively more diverse. It is dominated by grazing trails in prodelta, fine-grained sandstones with current ripple cross-lamination (Fig. 3). Sections traditionally considered of lacustrine origin, particularly in the eastern Paganzo Basin such as the Malanzán

Transgressions related to the demise of the Late Paleozoic Ice Age Karoo and Associated Basins (Paraná Basin, FalklandMalvinas Islands, and Sauce Grande–Ventana Fold Belt)

Formation (Fig. 1, M in Fig. 3), have been recently reinterpreted as having been connected to an inland sea. The mudstone-dominated section (Member 2 of the Malanzán Formation, Andreis et al., 1986) overlying the basal glacial diamictites and associated breccias and conglomerates (Member 1, Andreis et al., 1986) are now considered part of the postglacial transgression of glacial episode II owing to the discovery of brackish-water acritarchs, green algae that could tolerate a wide range of salinities like Botryococcus, and marine acritarchs (Navifusa and Greinervillites) (Gutiérrez and Limarino, 2001). This section represents the easternmost extent of the postglacial transgression in the Paganzo Basin (Limarino et al., 2002, 2006); the association of the acritarchs and green algae with abundant spores, log fragments, and coalified particles suggests proximity to the coast. Also the coexistence of specimens of marine salinity with other elements of brackish- to fresh-water habitats indicates drastic salinity fluctuations probably related to extreme fresh-water release from ice melting into an inland sea (Buatois et al., 2006, this volume) connected with an open sea to the west (CalingastaUspallata Basin).

+ +

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Karoo Basin The Karoo Basin (Figs. 1 and 6) developed as part of a series of linked foreland basins behind the Panthalassan margin (Veevers et al., 1994). Sedimentation was initially confined to the southern Karoo Basin adjacent to the Cape fold belt, and then spreading northward. The basin can be subdivided into three distinct segments: (1) a foredeep, close to the rising Cape fold belt, transgressed by an interior seaway; (2) a forebulge uplifted above base level; and (3) a shallow backbulge depozone to the north (Catuneanu, 2004). This subdivision broadly corresponds to the Dwyka glaciogenic facies distribution proposed by Visser (1986, 1989), with a platform facies association in the south and a valley facies association in the north. Only in the forebulge area restricted to the northernmost part of the basin, nonmarine deposits prevail with colluvial and glacial-outwash alluvial and fluvial facies (Faure et al., 1996; Bordy and Catuneanu, 2002). Lacustrine dropstone facies have been reported to the north,

~

West Falkland

Ind

B

~ ~ ~ ~ ~ ~ ~ ~~ ~ ~~ ~ ~ East ~ ~ ~ ~ ~ ~ ~ Falkland CZ 0 ~ 24° E 26° E 28° E CZ 22° E Drakensberg Group (Jurassic)

Lebombo Group

Stormberg Group

(Jurassic)

(Triassic - Early Jurassic)

Location of cross section B (modified from Visser, 1987a)

Beaufort Group (latest Permian-Triassic)

Ecca Group Distribution of Vryheid Formation

Dwyka Fm. + Ecca Gr. Dwyka Formation

32° S

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ia

n

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300 km

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~

(Ordovician-Devonian)

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Precambrian Basement

Location of borehole DP1

DP1 (from Faure and Cole, 1999)

Figure 6. Generalized map of the Karoo Basin with location of A–B cross section (Fig. 7), DP1 borehole, and areal extent of Vryheid (Lower Ecca Formation). Pre-breakup Falkland Islands position based on Marshall (1994).

12

López-Gamundí

suggesting deposition in glacial or periglacial lakes. Marine fossils have been reported from the base of the Prince Albert Formation (McLachlan and Anderson, 1973) in the central and western parts of the Karoo Basin. The Dwyka sediments exhibit a wide gamut of facies (massive to thin-bedded diamictites and dropstone (IRD) shales, “platform facies” of Visser, 1986, 1989) arranged in four deglaciation sequences (Fig. 7, Visser, 1997b); facies in the overlying Prince Albert Formation are dominated by siliceous shales with phosphatic chert and carbonate lenses and concretions. Glendonites, carbonate pseudomorphs after ikaite (CaCO3·6H2O), form at near freezing temperatures and have been reported from just above the top of the Dwyka deposits (McLachlan et al., 2001). Pioneering work by Visser (1989) on geochemical studies on carbon (C) and sulfur (S) contents shed some light on the depositional conditions of the Dwyka Group and the overlying Prince Albert sediments in the foredeep depozone (platform facies of Visser, 1986,

1989). Seawater is an abundant source of S in sulfate; in contrast the concentration of sulfate is much lower in fresh water (Berner and Raiswell, 1984). Four samples from the postglacial shales (Prince Albert Formation) indicate a C/S ratio of <2, which corresponds with marine mud deposition. Furthermore, the positive S intercept in the C versus S plots suggests euxinic bottom conditions with H2S-laden water being present (Leventhal, 1987). Mudstones within the glacial Dwyka section have a C/S ratio of 5 to 30, which suggests possible brackish conditions during deposition. Geochemical analyses based on rare-earth-element (REE) patterns, Sr concentrations, and 87Sr/86Sr ratios from calcite concretions found in the uppermost Dwyka sediments and overlying Prince Albert shales along the Cape fold belt suggest that fresh-water conditions prevailed (Herbert and Compton, 2007). This conclusion led to the inference of a lacustrine origin for this time interval in the entire basin, despite the most reliable paleogeographic reconstructions, which indicate the presence of

A

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North GRAAFF-REINERT

EDEN-BURG TROMPS-BURG

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Diamictite with sste. bodies

Stratified diamictite (Dms)

Diamictite (Dmm) with boulder beds

DS-2

DS-1

Borehole section

0

100 km

HS

Figure 7. North-south cross section A–B (see location in Fig. 6) across the Karoo Basin; datum at the base of Whitehill Formation (modified after Visser; 1997a). Deglaciation sequences (DS) and facies associations from Visser (1997b). Note that uppermost part of the DS-1 and DS-2 are represented by the dropstone shale facies association, herein interpreted as the early part of the transgressive systems tract (TST). In contrast, the TST of the DS-3 (Prince Albert Formation) shows both its early phase (dropstone shale facies association) and late phase (shales with subordinate fine-grained sandstones). Grain-size scale in logs: Si—silt; S—sand; G—gravel. HS—horizontal scale; VS—vertical scale.

Transgressions related to the demise of the Late Paleozoic Ice Age

A

S total (wt %) 1

0

3

2

-60 -65

Tieberg Formation

-70 -75

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Whitehill Formation

-85 -90 -95 -100

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a large inland sea or an interior seaway in a foreland-backarc setting (Von Brunn and Gravenor, 1983; Visser, 1989, 1991; Veevers et al., 1994; Johnson et al., 1996). The existence of a large inland sea with fluctuating salinity during Dwyka and Prince Albert times is also consistent with the apparent discrepancy of some geochemical results. Deglaciation is also indicated by an increase of organic matter content during a transgression. TOC values have been used to differentiate the onset of the deglaciation phase in the Karoo basin; TOC values up to 7% were reported for the lowermost part of the Prince Albert by Scheffler et al. (2006). The organicrich part of the postglacial transgression also was identified by Faure and Cole (1999) in the DP1 borehole (see Fig. 6 for location), with values up to 12% TOC (Fig. 8A). TOC versus S plots show a wide range of salinities ranging from marine to prevailing nonmarine conditions (Fig. 8B). High V/Cr ratios in concert with increasing TOC values indicate anoxic bottom-water conditions at the onset of the postglacial transgression (Scheffler et al., 2003, 2006). The accumulation of phosphatic sediments points to high bioproductivity resulting from elevated nutrient supply. The final retreat of the glaciers is recorded in the southern and eastern Karoo Basin by a concomitant rise of the chemical index of alteration (CIA) and Rb/K ratios (Scheffler et al., 2006). Predominantly full marine conditions prevailed, and the establishment of anoxic conditions is indicated by the occurrence of TOC-rich sediments containing mineral phases such as apatite and pyrite. Additional evidence in favor of marine conditions during the deglaciation phase is provided by the Rb/K ratio, used as a salinity proxy. Higher Rb/K ratios during the late phase of each deglaciation sequence in the Dwyka Group and during the lowermost Prince Albert suggest increasing salinity during those time intervals (Scheffler et al., 2006). Furthermore, the CIA and Rb/K high values suggest increased salinity during the last phase of Visser’s (1997b) deglaciation sequences (DSs) II, III, and IV (Scheffler et al., 2003). The northeastern part (the shallow backbulge depozone) of the Karoo Basin became ice free as a result of the collapse of the marine ice sheet. Based on studies along the northeastern margin of the basin and analogies with Pleistocene deglaciation models, Haldorsen et al. (2001) identified three glacial phases: an initial phase of fast (a few thousand years) marine ice retreat of 400 km over the northeast with glaciers grounded in the shallower areas around the shore of the basin, followed by a second phase characterized by smaller continental ice sheets that remained more or less stationary for several tens of thousands of years; during this second phase, massive glaciomarine muds with dropstones accumulated in the open basin; the last phase corresponds to final deglaciation, which might have taken 10–20 ka. According to Haldorsen et al. (2001) the entire Early Permian deglaciation of the northeast margin of the basin might have been completed within thousands rather than millions of years. The valley-glacier-retreat facies association is represented by the Vryheid Formation along the northern cratonic margins of the Karoo Basin, particularly on the northeasternmost corner

13

2 1 0 0

2

4

6

8

10

12 14

16

TOC (wt %) Figure 8. Geochemical characteristics of postglacial transgressions. (A) Total organic carbon (TOC) content (wt%) and sulfur (S) content (wt%) in borehole DP1 (Faure and Cole, 1999). See text for discussion. (B) TOC content (wt%) vs. S content (wt%) plot; the field at the lower right is considered fresh water, and the field at the upper left is considered marine. Values for tasmanite beds from Revill et al. (1994).

of the basin in Natal, Orange Free State, and Transvaal (Fig. 6), where the Vryheid Formation (Artinskian, Cairncross, 2001) forms a clastic wedge that pinches out basinward (Tankard et al., 1982). Coals of the Vryheid Formation rest directly on basement, Dwyka glaciogenic sediments, or marine postglacial shales (Pietermaritzburg Formation). Where present, the underlying Dwyka sediments contrast drastically with the basinward platform facies association of Visser (1986, 1989); the latter is characterized by massive and homogeneous diamictites (Dmm); the former association (the valley/inlet facies association, Visser, 1986) is represented by varied lithologies (with a higher participation of sandy matrix-supported conglomerates and associated sandstones) resting on an irregularly dissected paleosurface (Von Brunn, 1996).

14

López-Gamundí

The thickest (>5 m) coals in the Vryheid Formation are the lowest ones; they thin and pinch out against relatively steep-sided walls of paleovalleys scoured by glacial action (Cairncross, 1989). Cold-temperate Glossopteris-Gangamopteris flora is present in the floral assemblages of the lowermost peat (coal) swamps. Pulsating glacial retreat produced meltwater discharge that reworked unconsolidated till and introduced coarse, gravelly detritus. Thus, the lowermost coals are associated with conglomerates and sandstones deposited in braided plains by low sinuosity rivers (Le Blanc Smith and Eriksson, 1979; Cairncross, 1980). Some paleotopographic lows in the basement were sheltered from the coarse glaciofluvial influx, and lacustrine shales (rhythmites) and deltaic sediments capped by thin coals were deposited. These lowermost coals are capped by an extensive transgressive phase of shales and siltstones. Trace fossils (i.e., Skolithos and Planolites) described from cores in bioturbated levels between coal seams of the Vryheid Formation in Transvaal have been interpreted as part of abandonment phases in deltaic settings (Stanistreet et al., 1980). Glauconite pellets, interpreted as glauconitized fecal pellets, are associated with one of the bioturbated layers, suggesting a marine influence during the abandonment phase. According to Cairncross (1989), two factors related to the postglacial transgression could have contributed to the preservation of relatively thick coals at the base of the Vryheid Formation. Bacterial and microbial activity may have been suppressed by the cold temperatures, and the destruction of peat was not intense, allowing thicker coal-seam formation. Also, meltwater discharge probably promoted a high water table (cf. Bohacs and Sutter, 1997) that would, in turn, have maintained the swamps and enhanced peat preservation. Farther north in the Karoo-age rifts of Namibia, such as in the southeastern Kalahari Basin (9 in Fig. 1), anoxic conditions were established during the postglacial marine transgression, as evidenced in the lower Ecca shales (Tswane Formation). The organic-rich shales, which are carbonaceous in places, contain up to 50% TOC (Scheffler et al., 2006). The shales are underlain by fluvioglacial and glaciolacustrine deposits (Visser, 1997a). These deposits can be assigned to the valley-glacier-retreat facies association that dominates the updip portions of the glaciated basins. Farther basinward in the Huab and Kalahari Basins (8 and 9, respectively, in Fig. 1) the four deglaciation sequences (DSs) defined by Visser (1997b) for the main Karoo Basin have been identified (Bangert et al., 1999; Stollhofen et al., 2000). Maximum flooding intervals have been identified as IRD-free dark mudstones, locally known as the Ganigobis and Hardap Shale Members for DS-II and DS-III, respectively. The DS-II (with a maximum thickness of ~200 m) consists of a fining-upward succession of basal massive diamictites and outwash conglomerates that pass upward to rain-out diamictites and associated sandy gravity-flow deposits that grade into IRD-bearing mudstones. The succession is capped by the Ganigobis Shale Member. This unit is made up of dark, IRD-free mudstones deposited by offshore suspension settling, interbedded with scarce, thin beds of fine-grained sandstones interpreted as distal turbidites (Stollhofen

et al., 2000). Siliceous and phosphatic concretionary zones with fish remains and invertebrate fossils (gastropods, brachiopods, echinoids, crinoids, foraminifers, radiolarians), and fallout tuff beds are present in the Ganigobis Shale Member, where TOC contents between 0.5 and 1.4% and an S content of up to 1% have been reported (Grill in Stollhofen et al., 2000). A similar vertical facies stacking is present in DS-III ( ~130 m thick, Stollhofen et al., 2000) with rarely preserved massive diamictites at the base, followed by crudely bedded rain-out diamictites and massive to cross-bedded sandstones that grade up into IRD-rich sandy mudstones, and finally offshore IRD-free mudstones with bivalve shells of the Eurydesma fauna (Hardap Shale Member). The DS-IV is thinner (maximum thickness of 70 m) and made up, from base to top, of sediment gravity-flow sandstones followed by IRD-rich, sandy mudstones that grade upward into IRD-poor mudstones with thin interbeds of turbidite sandstone. Paraná Basin In the pre-drift neighboring Paraná Basin (Fig. 1) glaciogenic sediments of glacial episode III correspond stratigraphically to the Itararé Group. Gravenor and Rocha-Campos (1983) defined a marine-freshwater facies in the Itararé deposits. It consists of well-stratified thin-bedded diamictites (Dms) intercalated with sandstones and shales with IRD, massive diamictites in beds up to 1 m thick interbedded with shales, rhythmites, and mudstones, some with carbonate nodules. Marine fossils are present in the shales and mudstones. This facies is considered by Gravenor and Rocha-Campos (1983) as having been deposited in a basinal environment during deglaciation, and is capped by marine shales. At least four stratigraphically different basinwide marine “horizons” have been recognized in the basin (Rocha-Campos and Rösler, 1978), and they are probably similar to the final stage of the deglaciation sequences described for the Karoo Basin (cf. Fig. 7) by Visser (1997b). The basin margins show unequivocal evidence of glacier advances and retreats. During the late Early Permian, marine conditions expanded in the basin, and a final postglacial eustatic flooding is recognized by the widespread occurrence of relatively thin, fossiliferous marine beds in the uppermost part of the Itararé sequence (dos Santos et al., 1996). Three settings for the glacial-postglacial transition in the Paraná basin were defined by dos Santos et al. (1996). The terrestrial setting is characterized by subglacial diamictites with evidence of subglacial-style soft sediment deformation, massive or cross-bedded fluvial sandstones, and thin coal layers. The valley setting consists of diamictites as part of the basal fill of valleys controlled by basement fault zones and probably scoured and deepened by the ice, fluvial sandstones, massive or with cross bedding, and lacustrine mudstones and rythmites with IRD. In places this glacial terrestrial facies association is overlain by mudstones and siltstones with marine acritarchs and Tasmanites. Botryococcus is present in this postglacial section. The terrestrial and valley settings of dos Santos et al. (1996) broadly correspond to the valley glacier retreat facies association described in this contribution. The proximal glaciomarine setting consists of thin, discontinuous, subglacial,

Transgressions related to the demise of the Late Paleozoic Ice Age massive diamictites (Dmm), subglacial meltout-fluvioglacial coarse-grained conglomerates, debris flow, stratified diamictites (Dms[r]) and rhythmites with dropstones (França et al., 1996; Rocha-Campos et al., 2008). This facies association grades up into the overlying postglacial, transgressive marine shales of the Rio do Sul Formation. Thin, transitional sequences between the rhythmites and marine shales consist of shallow-marine, heterolithic sequences of siltstones and sandstones with lenticular bedding and intersecting wavy laminae indicative of wave action (Canuto, 1993, in dos Santos et al., 1996). Dropstones are common in the lower few meters of marine shales, but they gradually disappear upward. This transgressive facies association culminates the retrogradational pattern punctuated by short-term, transgressive-regressive episodes that characterize the deglaciation process (dos Santos et al., 1996). A glacial terrestrial setting is characterized by subglacial diamictites with evidence of subglacial-style soft sediment deformation, massive or cross-bedded fluvial sandstones, and thin coal layers. Along the basin margins the glacial facies are followed by littoral and deltaic sequences interpreted as basinward progradational wedges (dos Santos et al., 1996), where eventually paralic peat swamps developed. These progradational, commonly coal-bearing sequences have been identified along the northern (Martini and Rocha-Campos, 1991), western, and eastern margins (Buatois et al., 2006). Dos Santos et al. (1996) speculated about the correlation between these coal-bearing zones and the postglacial marine “horizons.” Lateral variability of sediment supply related to glacial activity along the basin margins during glacial and early postglacial times could be the principal factor accounting for the different number of sequences identified. However, the vertical arrangement of each sequence seems to be strikingly similar. Canuto et al. (2001) described seven fining- to coarsening-upward sequences for the Itararé Subgroup in their outcrop-based study along the eastern margin of the basin. These sequences range between 50 and 180 m in thickness; their base is characterized by an erosional surface (striated bedrock or intraformational striated pavements). Thin (>1 m), laterally discontinuous beds of compacted massive diamictites (Dmm), interpreted as lodgment tills, rest on the striated surfaces. This basal facies is followed by tabular beds of massive diamictites (Dmm), locally bedded (Dms) and interpreted as subglacial ablation tills associated with cohesive debris flow beds represented by massive to bedded diamictites (Dms[r]) with soft sediment deformation. This facies association is grouped in an LST by Canuto et al. (2001). The LST is followed by sandstones, siltstones, and claystones with dropstones (Fld) of a TST, which, in turn, is overlain by deltaic sandstones stacked in a progradational pattern (highstand and regressive glacio-isostatic systems tracts). The uppermost of these sequences (S7) represents the period of final deglaciation in the Paraná Basin. This glacial-postglacial transition lies stratigraphically between the uppermost part of the Itararé Group (Rio do Sul Formation), containing the LST and TST, and the basal Triunfo Member of the superjacent Rio Bonito

15

Formation (HST). A similar sequence stratigraphic framework was proposed by Vesely and Assine (2004) in their study that incorporated both outcrop and well-log data. They identified five depositional sequences bounded by disconformities traceable over 400 km across the basin in an E-W depositional strike section. Three successive facies associations were defined within each sequence, but the lower and upper ones are absent in places. The lower facies association occurs only in the two lowermost sequences and is made up of subglacial facies (mostly Dmm); it corresponds to a glacial-maximum systems tract as defined by Vesely and Assine (2004). This tract grades upward into conglomerates and sandstones, which in turn are overlain by diamictites, turbidites, and shales with dropstones, labeled by the authors as a deglaciation systems tract (equivalent to the transgressive systems tract of Canuto et al., 2001). The deglaciation facies rests directly on the bounding erosional surfaces where the subglacial facies of the glacial maximum system is absent basinward. The fine-grained laminated facies of the upper part of the deglaciation systems tract represents the record of maximum glacial retreat during interglacial periods. Highest gamma-ray readings were used to determine the stratigraphic position of the MFSs within the deglaciation systems tracts. Falkland-Malvinas Islands Similarities between the late Paleozoic glacial-postglacial sections in the Karoo Basin and the Falkland Islands (Fig. 9) were first suggested by Adie (1952). Later paleomagnetic (Mitchell et al., 1986; Taylor and Shaw, 1989), radiometric (Mussett and Taylor, 1994), and paleocurrent (Storey et al., 1999; Trewin et al., 2002) data further confirmed Adie’s original correlation. The pre-breakup reconstructed position of the Falkland microplate east of the coast of South Africa (Figs. 1 and 6) is the result of an eastward translation and a 180° rotation from its present location (Marshall, 1994). The glaciogenic deposits are grouped under the Lafonia Formation (Frakes and Crowell, 1967; Scasso and Mendía, 1985; Bellosi and Jalfin, 1987), and the overlying postglacial sediments are part of the Port Sussex Formation (Frakes and Crowell, 1967). The contact between the uppermost Lafonia diamictites and the lowermost member of the Port Sussex Formation (Hells Kitchen Member) was characterized as an “abrupt undulating surface with a local relief of 0.25 m” by Trewin et al. (2002), who also note that the member thickness ranges from 3.5 to a maximum of 10 m and is made up of a coarsening-up cycle grading from “black fissile mudstone to fine- and medium-grained sandstones with granules and pebbles deposited as dropstones” (p. 9). Trewin et al. (2002) interpret the contact between the Lafonian diamictite and the base of the Hells Kitchen Member as an erosional event that probably developed under shallow-water conditions during a transgression related to deglaciation. This contact is considered herein as a transgressive ravinement surface (TRS) that initiated the postglacial TST. Trewin et al. (2002, p. 9) suggest that the Hells Kitchen Member sediments were deposited “probably during a transgression associated with deglaciation” and compared the section with a similar

16

López-Gamundí Si

SG Si

Bonete Fm.

SG

Ripon Si S

G

Fm.

Piedra Azul Fm.

Collingham Fm. Whitehill Fm.

Prince Albert Fm.

Sauce Grande Fm.

Dwyka Fm.

Highstand systems tract Transgressive systems tract

Port Sussex Fm.

Lowstand

Lafonia Fm.

systems tract

100 m 0 Black shales (Fl)

Massive diamictite (Dmm)

Shales (Fl) and fine-grained sandstones (Sr)

Coarse sandstone

Dropstone shale (Fld), thinly bedded diamictite (Dms) and shales (Fl)

Fine to medium sandstone

Stratified diamictite (Dms)

Basement

Figure 9. Correlation of the postglacial mudstones (glacial epsiode III) across the Ventana fold belt (V), Karoo Basin (K), and Falkland Islands (FI). Grain-size scale in logs: Si—silt; S—sand; G—gravel.

facies arrangement described by Visser (1993) for the base of the Prince Albert Formation in the Karoo Basin. The Hells Kitchen Member is overlain by interbedded black, pyritic, partly carbonaceous shales and siliceous mudstones (Black Rock Member, with a minimum thickness of 120 m); TOC content in the carbonaceous shales ranges ~2%–3% in the east to exceptionally up to 40% toward the west (Trewin et al., 2002). TOC versus S plots for two samples of the Black Rock Member indicate fresh-water salinities (Fig. 8B). Sauce Grande Basin–Ventana Fold Belt Keidel (1916) first recognized the similarities between the Cape fold belt and the adjacent Karoo Basin in South Africa, and the Ventana fold belt and contiguous Sauce Grande Basin in Argentina (Fig. 9). The similarities are particularly striking for late Paleozoic times, when the overall basin evolution and subsidence histories in both regions are considered (Du Toit, 1927; López-Gamundí and Rossello, 1998). Sedimentation in the Sauce Grande Basin adjacent to the Ventana fold belt began with

glaciogenic deposits by the Late Carboniferous (Keidel, 1916; Harrington, 1947; Coates, 1969; Amos and López-Gamundí, 1981b). This is approximately the same time interval when sedimentation, dominated by glacial deposits as well, began also in the Chaco-Paraná and Paraná Basins (Fig. 1). The basal fill of the Gondwana cycle in both basins is characterized by glacial deposits, mostly nonmarine in the Chaco-Paraná Basin (Russo et al., 1987; Fernández Garrasino, 1996; Winn and Steinmetz, 1998) and predominantly marine in the Paraná Basin (Rocha-Campos and dos Santos, 1981; Zalán et al., 1990; França and Potter, 1991; França, 1994; Rocha-Campos et al., 2008). Late Paleozoic glacial deposits exposed in the Ventana fold belt are grouped under the Sauce Grande Formation. Equivalent diamictites and overlying finer grained units (Piedra Azul and Bonete Formations) also have been reported from offshore wells (Lesta et al., 1980; Amos and López-Gamundí, 1981b; Fryklund et al., 1996; Juan et al., 1996). The age of the Sauce Grande Formation is poorly constrained as pre–Early Permian by the presence of Eurydesma fauna and Glossopteris flora in the Bonete

Transgressions related to the demise of the Late Paleozoic Ice Age Formation and post–mid-Carboniferous on the basis of palynological studies in equivalent sediments found in offshore wells (Archangelsky, 1996). The base of the Sauce Grande Formation lies with regional unconformity on Devonian metasedimentary basement. The unit is mostly made up of massive to crudely stratified diamictites with subordinate rhythmites, sandstones with ripples, and scarce conglomerates toward the top (Andreis et al., 1987). The massive (Dmm) and stratified (Dms) diamictites have been interpreted as glaciomarine (rain-out tills) partially remobilized downslope by gravity flows (Dms[r]) (Coates, 1969; Andreis, 1984; Andreis and Torres Ribeiro, 2003). The marine setting is confirmed by the presence of a solitary marine bivalve in the diamictites (Harrington, 1955). Detailed sedimentological studies by Andreis and Torres Ribeiro (2003) allowed subdivision of the Sauce Grande Formation in three megacycles. The lower megacycle (maximum thickness, 700 m) is composed of abundant diamictites, sandstones, and scarce conglomerates. The middle megacycle (~50 m thick) contains only sandstones and conglomerates. The upper cycle (~350 m thick) is made up of abundant fine- to coarse-grained sandstones; thick-bedded, massive, clast-poor, muddy diamictites (Dmm); thin-bedded stratified diamictites (Dms[r]); and shales, with scattered dropstones. The Dmm facies is interpreted as rainout tills, whereas the Dms(r) facies is considered glacial material remobilized downslope as gravity (mostly debris) flows. Andreis and Torres Ribeiro (2003) interpret the finer grained nature (evidenced by increasing participation of shales and sandstones to the detriment of diamictites, particularly of the clast-rich subtype) of the uppermost megacycle as a response to the transgression caused by glacier retreat that continued during Piedra Azul times. Dickins (1984) also relates the Piedra Azul transgression to the early Sakmarian glacioeustatic sea-level rise. The facies types and their vertical stacking in the Sauce Grande upper megacycle show similarities to the deglaciation sequences described by Visser (1997b) for the Karoo Basin (cf. Figs. 7 and 9). The Sauce Grande sediments grade upward to shales, bioturbated mudstones with gastropods, and subordinate fine-grained sandstones with wave and current ripples (Piedra Azul Formation), passing upward into bioturbated mudstones with fossils of the Eurydesma fauna (Harrington, 1955; Rocha-Campos and Carvalho, 1975; Amos, 1980), and fine-grained sandstones with wave ripples and cross-bedding (Andreis et al., 1979). Coarsening upward parasequences (up to 40 m thick each) are common in the Piedra Azul Formation; they consist, from bottom to top, of mudstones, fine-grained sandstones with wave ripples, and channelized medium- to coarse-grained sandstones with trough crossstratification (Andreis and Japas, 1996). Gondwana Basins of Peninsular India The Gondwana basins of Peninsular India (Fig. 1) have been considered fault bounded troughs developed along preexisting zones of weakness imparted by Precambrian structural fabrics (Naqvi et al., 1974). The basin fill is commonly asymmetric, with an overall increase in thickness toward one of the boundary faults,

17

a feature typical of extensional rift basins. A significant strikeslip component has been proposed for these basins (Chakraborty et al., 2003; Chakraborty and Ghosh, 2005). The Talchir Formation in the Indian peninsular basins (Fig. 10) shows evidence of sedimentation under glacial influence (Ghosh and Mitra, 1975; Casshyap and Srivastava, 1987; Veevers and Tewari, 1995). Key evidence of glacial influence during Talchir sedimentation is provided by striated pavements on basement rocks and boulder pavements, striated and bulletshaped boulders in diamictites, and dropstone shales (Banerjee, 1966; Ahmad, 1975) Marine faunas have been reported at or near the top of the glaciogenic Talchir Formation in several Gondwanan basins of India (see Veevers and Tewari, 1995, for a review). These findings may be related to a marine transgression, interpreted by several authors as having been caused by a eustatic rise in sea level owing to deglaciation. Previous interpretations, however, originally considered that the glacial sediments and finer grained postglacial sediments were deposited at the margin of lake basins, despite the rather gradational passage to marine deposits with abundant invertebrate fauna dominated by bivalves in the Rewa Basin (Ghosh, 1954), Daltonganj-Rajhara Basin (Dutt, 1965, in Goswami, 2008), and Bokaro Basin (Sengupta et al., 1999) (Fig. 10A). The most distinctive element of these marine faunas is Eurydesma, a cold-water bivalve genus widespread in the Early Permian of the Gondwana Supercontinent (Harrington, 1955; Dickins, 1961; Dickins and Shah, 1977; Runnegar, 1979). An inspection of the section described by Gosh (2003) in the Satpura Basin (Fig. 10B) suggests that the black shales and subordinate, decimeter-scale limestone beds that rest on the glaciogenic Talchir diamictites can be interpreted as part of the postglacial marine transgression rather than as the result of sedimentation of predominately fine-grained (clay and silt) material and carbonate muds in a basinal lacustrine environment with a sporadic marine incursion, containing a marine fauna characterized by Eurydesma and presence of other bivalves, ostracodes, and foraminifers. Marine acritachs have been identified in several basins and, in a few cases, associated with Eurydesma (Venkatachala and Tiwari, 1988). Associated plant remains also have been reported in several localities (Gosh, 2003). The postglacial section in the Satpura Basin is a marine, deep-water facies association; isolated ripple trains can be interpreted as starved ripples produced by weak bi- or unidirectional bottom currents. This finding extends also this marine embayment of the Indian Peninsula farther west from the areas where marine faunas were originally described (Rewa, Daltonganj-Rajhara, Bokaro Sub-basins of the Damodar Basin, Fig. 10A). The alternative interpretation should invoke a setting in which deep lacustrine facies are systematically interrupted by open-marine shales with Eurydesma fauna deposited below wave base without any intervening facies. Bose at al. (1992) studied the sedimentary features of the Talchir sandstone facies in the West Bokaro Sub-basin (Fig. 10A) and interpreted them as the product of coarse-clastic sedimentation “in fjord-like glacier-fed coastal troughs” (p. 95). Supporting evidence for a fjord-like setting with low salinity from a fresh-water

18

López-Gamundí Micritic limestones

78°E

Son

N

86°E Daltonganj-Rajhara

Rewa Basin

24°N

Satpura Basin

B

∗

Fine-grained sandstones

Raipur

Starved ripples Burrows

Paleocurrents Marine fauna (Eurydesma)

Talchir Basin

Irai

Trough cross-stratification

Massive diamictites (Dmm)

Raniganj

Mahanadi Basin Bedasar

20°N

Jahria

B

Rahmajal

Damodar Basin

Son Basin

∗

∗ ∗

∗

West Bokaro

Hummocky cross-stratification

Shales / mudstones

Barakar Fm.

Bay

TST

of

Godavari Basin

A

100 km

∗

Localities with marine fauna

20 m

Bengal

Talchir Fm. Si S

G

Si S

G

Figure 10. (A) Gondwana deposits in Peninsular India. Note presence of marine fossils (Eurydesma fauna) in postglacial mudstones (glacial episode III). (B) Section of the glacial-postglacial transition in the Satpura Basin. Modified from Ghosh (2003). Grain-size scale in logs: Si—silt; S—sand; G—gravel.

contribution from retreating glaciers comes also from ichnological studies. Bhattacharya and Bhattacharya (2007) studied the ichnofacies associated with these postglacial intervals in the Raniganj Sub-basin (Fig. 10A) and concluded, on the basis of low ichnodiversity, sporadic distribution of the traces, small burrow dimensions, absence of any body fossils, and dominance of worms and annelids as trace-makers, that stressed environmental conditions (cold climate and low marine salinity) prevailed. These conditions are assigned to the influx of glacier meltout fresh water in an ice-marginal sea during climatic amelioration and deglaciation. The postglacial TST exhibits a deepening-upward facies trend initiated with (1) mudstones with dropstones; (2) interbedded laminated to massive siltstones, mudstones, and fine-grained sandstones with current and combined-flow ripples and hummocky cross-stratification; and (3) overlying IRD-free, massive, locally bioturbated mudstones. Nereites and Zoophycos ichnoassemblages are common in the two latter facies (Bhattacharya and Bhattacharya, 2007). These facies with abundant trace fossils are conspicuously IRD-free and indicative of a retreat of the icegrounding line with a concomitant influx of glacier meltwater into the basin and subsequent lowering of the marine salinity. The facies of siltstones, mudstones, and fine-grained sandstones is interpreted as the product of sedimentation in an open shelf influenced by waves, whereas the overlying mudstone facies represents offshore background sedimentation from suspension of clay-silt material. It is here proposed that the early TST is represented by the facies of mudstones with dropstones; the MFI

in the late TST is represented by the mudstone facies. A similar interpretation has been proposed for the upper Talchir interval in the adjacent West Bokaro Sub-basin (Fig. 10A) by Bhattacharya et al. (2005). The basal Talchir glaciogenic section is made up of breccias, matrix- and clast-supported conglomerates, and coarse-grained sandstones (their conglomerate-sandstone, TCS, facies association) and is followed upward by fine-grained sandstones with hummocky cross-stratification, sandy siltstones with dropstones, and subordinate clast-supported conglomerates (sandstone-siltstone, TSS, facies association). This fining-upward, retrogradational section culminates with alternating thinly bedded fine-grained sandstones and mudstones with marine fossils (mollusks) and thick, multistoried, fine-grained sandstones with common hummocky and swaley cross-stratification (fine sandstone–mudstone, TSM, facies association). Bhattacharya et al. (2005) interpret this retrogradational stacking pattern as the result of progressive deglaciation during a eustatic sea-level rise and deposition of shelf sediments under a transgressive phase for the TSS and TSM facies associations. Geochemical evidence is also supportive of an environment with significant fresh-water contribution. Bhattacharya et al. (2002) studied the geochemistry of calcareous nodules in finegrained sediments (siltstones, rhythmites) that rest on the coarse glaciogenic basal part of the Talchir Formation in the Damodar, Mahanadi, and Godavari Basins (Fig. 10A). Four nodules of similar composition (micritic) from the uppermost Dwyka tillite of the Karroo Basin were also analyzed for comparative purposes.

Transgressions related to the demise of the Late Paleozoic Ice Age The oxygen and carbon isotopic ratio (δ18O and δ13C) values from both Talchir and Dwyka samples were similar and indicate a fresh-water environment of formation. Tasmania Basin During the late Paleozoic, Tasmania (Fig. 1) was positioned close to polar latitudes ~70° (Late Carboniferous) and ~80° (Early Permian) (Scotese and Langford, 1995; Li and Powell, 2001). The Tasmania Basin (Fig. 1) was glaciated throughout the Late Pennsylvanian (Clarke and Forsyth, 1989; Dickins, 1996). Ice-flow indicators suggest ice centers on the west and northwest and depocenters on the east. Local topographic highs created a fragmented shelf where glacially influenced sedimentation took place (Clarke and Forsyth, 1989; Hand, 1993). There is consensus (Banks, 1981; Clarke and Forsyth, 1989; Hand, 1993) that the glaciomarine diamictite and associated rhythmites of the Tasmania Basin were deposited in a fjord-like seaway characterized by rain-out and settling of fines with minor coarse debris deposited by rafting and turbidity currents originating from the glacier grounding line (cf. Bartek and Anderson, 1991). The glaciomarine sediments pass upward to marine siltstones and mudstones (Woody Inglis Siltstone, Quamby Mudstone, and equivalent units) with abundant glendonite concretions and elements of the Eurydesma fauna (Clarke and Banks, 1975). This interval has been assigned to a widespread marine transgression that covered most of Tasmania as glaciers retreated (Hand, 1993) during the early–middle Sakmarian (Brakel and Totterdell, 1993). These fine-grained deposits are overlain by the Bundella Formation (late Sakmarian), which includes the Darlington Limestone. Dropstones and glendonite concretions are abundant within the limestone (Rogala et al., 2007); a high-abundance, low-diversity Eurydesma fauna of calcareous invertebrates (mostly brachiopods, bryozoans, and Eurydesma bivalves) has been identified (Clarke and Forsyth, 1989; Rogala et al., 2007). The limestones consist of bioclastic floatstones, rudstones, and grainstones deposited in neritic shelfal environments during sea-level highstands (Rogala et al., 2007). Cold conditions persisted until the Late Permian, as evidenced by dropstones and cold-water Eurydesma fauna in Kungurian to Capitanian beds (Clarke and Forsyth, 1989). The glacial-postglacial transition in the Tasmania Basin has been studied in outcrops and in subsurface (Figs. 11–13). Its contact is in general sharp (mudstones resting on mostly massive diamictites), but in a few localities a distinctive facies between the glaciogenic diamictites and the postglacial mudstones has been identified. This facies consists of thin (0.5–1 m of bed thickness), stratified diamictites (Dms) with dropstone clusters and pebble nests (Domack et al., 1993). These thin diamictite beds have flat, nonerosive lower contacts, and alternate with the isolated pebbly mudstones and dark gray to black mudstones; eventually the mudstones predominate. The postglacial black mudstones contain scattered dropstones; framboidal pyritic and large (up to 10 cm) calcareous concretions (glendonites) also have been identified (Domack et al., 1993). Glendonites are calcite pseudomorphs after ikaite, a hydrated type of calcium car-

19

bonate, CaCO3 · 6H2O (Suess et al., 1982), and seem to occur in those mudstone zones rich in organic matter derived from the unicellular alga Tasmanites. This mineral is suggested to be an authigenic precipitate, forming at low temperatures from interstitial waters of organic-rich sediments, undergoing microbial degradation, accumulating rapidly in cold bottom waters (Suess et al., 1982; Shearman and Smith, 1985). High alkalinity from decomposing organic matter also enhanced precipitation of ikaite (Bischoff et al., 1993; DeLurio and Frakes, 1999). Glendonites in the Sydney Basin (Fig. 1) are also commonly associated with organic-rich shelf facies with dropstones (Thomas et al., 2005; Selleck et al., 2007). The layers rich in Tasmanites (named tasmanite beds) are laminated and occupy a consistent stratigraphic position ~10– 30 m above the contact between the mudstone-prone postglacial interval and the underlying diamictic section (Figs. 11 and 12, Domack et al., 1993). Pyrite nodules and dropstones are abundant in the tasmanite beds. TOC values for the tasmanite levels in two boreholes (Douglas River and Ross 1) reach exceptionally ~20%, with high hydrogen indexes (HI = 868 mg HC/g TOC); even higher HI values >900 mg HC/g TOC have been reported from other localities (Fig. 14, Revill et al., 1994). These high HI values indicate that the kerogen contains hydrogen-rich type I organic matter (Tissot and Welte, 1984). Sulfur values obtained rarely exceed 2%; when plotted against TOC content they fall on both fields (marine sediments and fresh-water sediments), indicating conditions of variable salinity (Fig. 8B). As indicated by Domack et al. (1993), the association of glendonite with the TOC-rich tasmanite levels suggests a marine origin and also low temperatures (close to freezing), a condition required for the formation of the precursor of glendonite (ikaite) (Shearman and Smith, 1985). Rao and Green (1982) estimated independently a sea surface temperature for Early Permian Tasmania of –1.8 °C, close to the present average near the Antarctic ice shelf of –1.9 °C. Cold water deposition is consistent with the low diversity of the associated Eurydesma fauna (Clarke and Banks, 1975). The organic-rich tasmanite beds in the Woody Island Formation and equivalent units represent condensed sections within an MFI; a similar view was proposed by Rogala et al. (2007). DISCUSSION Framework of the Postglacial TSTs Punctuation of long-term transgressions by repeated shortterm regressions is caused by the tendency of sediment supply rates to outmatch, for short periods, accommodation increase rates (Cattaneo and Steel, 2003). This is a particularly common pattern in most transgressions during periods devoid of glaciations. This punctuated shoreline movement has a complex zigzag shoreline trajectory (Helland-Hansen and Gjelberg, 1994) that results in a retrogradational (or backstepping) stacking pattern (retrogradational parasequence set, Van Wagoner et al., 1990). As part of the Pleistocene Ice Age, the late Quaternary transgressive

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Figure 11. Cross section showing stratigraphic position of tasmanite levels within the postglacial mudstone section in the Tasmania Basin. From Domack et al. (1993).

deposits may represent an unusual record of high-frequency and high-amplitude sea-level oscillations driven by glacioeustasy, and therefore better analogues for the sections previously discussed. Rather, the postglacial transgressions analyzed herein seem to be related to abrupt upward-deepening of facies and the scarcity of surfaces of ravinement (erosion by wave action), culminating in a level of deepest facies, commonly termed the maximum flooding interval (MFI) or surface (MFS). Direct juxtaposition of offshore mudstones over glacial deposits (most commonly rain-out tills or less abundant subglacial tills) is not unusual. In other cases a diagnostic intervening facies association dominated by thin-bedded debris flow deposits with IRD and dropstone shales is present. This facies association is the result of the complex interaction between glacier-derived sedimentation, gravity (debris flow) currents, and rain-out deposition. Although

variations within this theme are possible along a continental margin subjected to deglaciation, as sediment supply may be laterally variable in a basin, the rate of sea-level rise is significantly higher than any other process. In that sense the postglacial TST identified in the late Paleozoic Gondwanan basins can be equated to the late Quaternary transgressive deposits resulting from an unusual record of high-frequency, high-amplitude sea-level rise driven by glacioeustasy. The response to this type of postglacial retreat is the creation of a continuous transgression triggered by steady (and fast) sea-level rise (Curray, 1964; Cattaneo and Steel, 2003). Experimental sequence stratigraphy provides insights into some of the salient features of fast transgressions. Heller et al. (2001) showed that a rapid base-level-rise rate leads to an abrupt shoreline retreat with maximum flooding at the end of the maximum base-level rise. They also noticed that the magnitude of

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Figure 12. Douglas River borehole section (modified from Domack et al., 1993), with TOC (total organic carbon) values for the Tasmanites-rich interval. Grain-size scale in log: Si—silt; S—sand; G—gravel. See location in Figure 11.

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transgression is far greater for the rapid base-level rise than during the slow rise, and that the time of maximum flooding is nearly synchronous with the end of the rapid base-level rise. In cases with slow rises, maximum flooding takes place far earlier than at the end of the rise.

B

Figure 13. Cores from the glacial-postglacial transition in the Tasmania Basin; see Figure 11 for location. (A) Massive diamictite (Dmm), probably of rain-out origin at base of Ross-1 well. (B) IRD-dominated section, Eaglehawk. (C) Stratified diamictites (Dms) with soft sediment deformation, Eaglehawk. (D) Organic-rich shale, Tasmanites-rich interval, Douglas River. Bar scale: 10 cm. IRD—ice-rafted debris.

D

Even in the cases where postglacial offshore mudstones are directly resting over the glacially derived, diamictitic section, the interval with the highest TOC values is invariably slightly higher in the interval dominated by offshore mudstones; thus most maximum flooding shales analyzed herein do not seem to be basal

Transgressions related to the demise of the Late Paleozoic Ice Age 1,000

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Figure 14. Geochemical characteristics of the tasmanite beds. From Domack et al. (1993).

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transgressive black shales sensu stricto (BT model of Wignall, 1991, 1994). The organic-rich, condensed sections in the Falkland Islands and Tasmania, a few meters above the contact with the diamictitic succession, are examples of this type of maximumflooding black shales. In other, leaner postglacial mudstone sections the maximum flooding interval is inferred by the abundance of marine fossils and dominance of fine-grained sediments (cf. most sections in the Calingasta-Uspallata and Paganzo Basins, the Piedra Azul interval in the Ventana fold belt, and the Lower Barakar Formation in several basins of Peninsular India). Postglacial TSTs through Space and Time As noted by Andrews (1997) the stratigraphic record may not be able to furnish the nuances of an unstable ice sheet at the time of its disintegration, especially without precise chronological control such as in the case of the LPIA. Thus, although appealing, it is somewhat incorrect to infer direct correlations between glacial histories and ice volume record. Despite this potential impossibility, the stratigraphic record of the late Paleozoic glacial-postglacial transition shows at hierarchies equivalent to third-order cycles a remarkable consistency in terms of facies associations, facies stacking patterns, and sequence-stratigraphic framework evolution. Other Occurrences of Postglacial TSTs Related to the LPIA The common characteristics of the glacial-postglacial transition highlighted in this contribution for the Calingasta-Uspallata and Paganzo Basins in western Argentina, the Karoo Basin in South Africa, the Gondwana basins of India, and the Tasmania

Basin in Australia seem to be present in other basins of the Gondwana Supercontinent affected by the LPIA. In Oman (Fig. 1) the glaciogenic sequence (Lower Permian Al Khlata Formation) is developed in a rift setting. It rests disconformably over Proterozoic basement and consists of diamictites, conglomerates, sandstones, and mudrocks ranging from glaciofluvial, glaciolacustrine, and alluvial to paralic environments (Levell et al., 1988; Al-Belushi et al., 1996; Martin et al., 2008). Evidence of glacial abrasion is provided in outcrops by striated surfaces on the dolomitic basement of late Proterozoic age (Braakman et al., 1982). A paleo– ice flow from northeast to southwest was inferred by Al-Belushi et al. (1996). Southward in the subsurface of the South Oman Salt Basin the Al Khlata diamictites were entirely deposited by rain-out and debris flow with no evidence for the preservation of true tillites (Aitken et al., 2004). The upper part of the Al Khlata Formation consists of a sheetlike diamictite abruptly overlain by the Sakmarian Rahab shale, which includes varve-like laminated mudrocks (rhythmites) with dropstones deposited in a large fresh-water to brackish-water body, according to Hughes Clarke (1988) and Levell et al. (1988). The dropstones progressively disappear upward, leading to Levell et al. (1988) to interpret this uppermost section of the Rahab shale as a deglaciation sequence. This unit is followed by the Saiwan Formation. In subsurface the Saiwan Formation unconformably overlies fine-grained sandstones and siltstones of the Rahab shale or rests directly on diamictites of the Al Khlata Formation. Its lowest part (lower member of the Gharif Formation) includes a restricted marine interval with the acritarchs, termed the maximum flooding shale by Guit et al. (1995) and interpreted as a possible postglacial eustatic flooding event (Stephenson and Osterloff, 2002). The base

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of the Saiwan Formation, dated biostratigraphically as late Sakmarian (Angiolini et al., 2003), includes bioclastic sandstones with assemblages of filter-feeding brachiopods and bryozoans, which compare closely with the postglacial depositional sediments described from low-energy species of Antarctic shelves and indicate, according to Angiolini et al. (2003), the final stage of the Gondwanan deglaciation. TSTs Associated with the Late Ordovician Glaciation The proposed correspondence between lowstand incision of paleovalleys filled with glacial sediments during a time of ice buildup and subsequent glacioeustatically induced transgression immediately after glacier retreat is not exclusive to the LPIA. The Late Ordovician glaciation, present mainly in intracratonic basins of Mauritania, Mali, Morocco, Algeria, Libya, Tunisia, Jordan, and Saudi Arabia (Fig. 15), is particularly abundant in examples of these mechanisms, particularly for the postglacial transgression and its well-known product in North African and Arabian Platforms: the world-class source rock informally known as “the hot black shales” (Keeley and Massoud, 1998; Jones and Stump, 1999; Lüning et al., 2000; Carr, 2002). The hot shales are defined by an arbitrary cutoff (>200 API units) in the gamma-ray curves of well logs (Lüning et al., 2000). This value correlates approximately with TOC values of 3% for maturities around the oil window. Initial marine transgression is marked by an erosional ravinement surface locally overlain by thin, residual shallow-marine sands (Boote et al., 1998). Recent paleoglaciological reconstructions of the Late Ordovician Saharan ice sheet (Le Heron and Craig, 2008) suggest a stepwise, southward recession from the ice maximum, followed by a postglacial transgression. The hot shales (basal Tanezzuft Shale) in the North African Platform were deposited during the initial transgression in paleodepressions (formed by

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previous glacial outwash valleys and structural depressions) that enhanced restricted circulation; stratigraphically they correspond to the early TST (Lüning et al., 2000). The postglacial shale unit receives different lithostratigraphic denominations (Aïn Deliouine Formation in Morocco, Argiles à Graptolites in Algeria, basal Tanezuft Shale in Lybia, Qusaiba Member in Saudi Arabia, Sahmah Formation in Oman, Batra Mudstone and Mudawwara in Jordan, Abba Formation in Syria, Akkas in Iraq, Dadas Formation in southeastern Turkey, and Ghakum Formation in Iran), but it seems to correspond to a single, fast contemporaneous episode of postglacial inundation across much of the North African and Arabian Platforms. Detailed studies of the associated graptolite faunas indicate that the deposition of the organic-rich hot shales was a synchronous event across North Africa through Arabia of Llandovery (Rhuddanian) age during the Early Silurian (Lüning et al., 2000; Miller and Melvin, 2005). In Arabia (1 in Fig. 15) the hot shales are highly fossiliferous and pyritic. They correspond to the basal part of the Qusaiba Member (Qalibah Formation), the principal source rock for Paleozoic hydrocarbons in Saudi Arabia (Mahmoud et al., 1992), and were deposited as a condensed sequence on a sedimentstarved shelf; the high TOC values have been interpreted as the result of high productivity in high-latitude water masses (Jones and Stump, 1999). Sequence stratigraphic studies (Lüning et al., 2000; Dardour et al., 2004) frame the record of this glacial event and its associated subsequent transgression in the context of an LST of glacial sediments covered by transgressive marine shales (TST). Dardour et al. (2004) proposed a Late Ordovician–Silurian (second-order) super-sequence comprising a periglacial lowstand, the Tanezzuft transgressive to early highstand shales, and the Acacus Formation

Illizi Basin

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Figure 15. Map of North Africa and Middle East regions, with record of Late Ordovician glacial deposits and associated Early Silurian postglacial transgressive deposits. Based on compilation from several sources by Le Heron et al. (2009). Numbers refer to outcrop and subsurface records of postglacial transgressions discussed in the text.

Transgressions related to the demise of the Late Paleozoic Ice Age late HSTs. The lower boundary for this super-sequence is well preserved in the northern part of the Murzuk Basin (SW Lybia, 2 in Fig. 15) and in the contiguous Illizi Basin (SE Algeria, 3 in Fig. 15) as deep, incised valleys. This topography was gradually infilled by a heterogeneous glacial (mostly glaciofluvial) lowstand clastic facies that grades laterally into a more uniform, distal facies in the Ghadames Basin (eastern Algeria, southern Tunisia, and NW Libya) farther north (4 in Fig. 15, Dardour et al., 2004). Several higher frequency glacial sequences are recognized that may reflect several cycles of glacial advance and retreat. The North African Platform was subsequently flooded by a transgressive facies, the Tanezzuft Shale, which contains at its base a short-lived Rhuddanian anoxic event represented by thin but regionally extensive organic-rich hot shales (Lüning et al., 2000), interpreted here as the MFI of the sequence. The aggregate thickness of the hot shales rarely exceeds 20 m in a thin (~40 m), hot-shale–bearing basal section of the Tanezzuft Shale in the Ghadames and Illizi Basins (Lüning et al., 2000). The postglacial TST was followed by regressive highstand sedimentation commencing in the late Llandoverian; the HST continued to the end of the Silurian with sedimentation of a northerly prograding shelfal to fluviodeltaic wedge. Similar scenarios have been described elsewhere for the same Late Ordovician–earliest Silurian glacial-postglacial transition. Turner et al. (2005) described lowstand channel incisions and fill of those channels with glaciofluvial and shoreface sandstones in southern Jordan (5 in Fig. 15) related to a fall of relative sea level during ice buildup in southern Arabia. Each incision has been correlated with the first glacial advance during a stage of reduced accommodation space. This stage was followed by glacial melting and marine transgressive filling of the incised valleys as accommodation space increased. This process was repeated four times with glacier re-advances evidenced by glacially scoured, striated surfaces. The final transgressive filling (equivalent to the postglacial TST here) is characterized by the hot blackshale interval locally known as the lower Batra mudstone. The base of the black shale is coincident with the MFS (Armstrong et al., 2005). Recently Armstrong et al. (2009) concluded that the Batra Formation black shale was deposited in a short-lived, permanently stratified marine basin where an influx of fresh water from melting ice and nutrients resulted in enhanced photic-zone primary productivity and organic matter sedimentation. They proposed the stratified basins and fjords of east Antarctica as possible modern analogues. Similar valleys incised to depths exceeding 600 m (McClure, 1978; Vaslet, 1990), and subsequently filled by glacial diamictites and pro-glacial sandstones, have been traced into the subsurface of northern Saudi Arabia with seismic data (McGillivray and Husseini, 1992). After the retreat of glaciers the Arabian Platform was flooded by a rapid, glacioeustatic sea-level rise (Konert et al., 2001) that set anoxic conditions for the deposition and preservation of organic-rich hot shales (Qusaiba shale). The postglacial transgression (TST) was followed by a thick (>1000 m) coarsening-upward sequence of shales and sandstones (HST) of

25

Llandovery to Pridoli age that prograded basinward (Mahmoud et al., 1992; Konert et al., 2001). A similar sequence is described for the glacial-postglacial transition by Le Heron et al. (2007) in the Anti-Atlas of Morocco (6 in Fig. 15) where stratified, clast-poor, sandy diamictites, lying directly above a striated surface, pass vertically into transgressive tidal deposits (sigmoidally cross-bedded sandstones), which in turn pass upward into offshore mudstones (Aïn Deliouine Formation). Rare outsize boulders in clast-poor diamictites deform underlying laminae, indicative of dropstones from icebergs. The overall fining-upward section has been interpreted as having been deposited during ice sheet retreat in the earliest Silurian. Asymmetry in Sea-Level Rates: Its Impact on the Stratigraphic Record The asymmetry (fast rates of sea level rise, cf. Fig. 2) in the eustatic sea-level curve defines a clear stratigraphic signature, especially in cases where the sediment supply rate is relatively moderate to low with respect to the sea-level rise. The extent of erosion and creation of a TRS depends on the rate of rise of relative sea level. In cases such as the postglacial TST with a relatively rapid sea-level rise, erosion is minimized. Thus TRSs are infrequent, and transgressive sediments are preserved. The postglacial signature is characterized by retrogradational stacking patterns. Two subtypes of TSTs, which should be considered end members within a continuum for the submarine-retreat facies association (Fig. 16), are identified: 1. Complete TSTs: In cases where the sediment supply rate is fast enough to keep up with the sea-level rise rate, the lower part of the TST is dominated by rain-out processes, gravity flows (mostly proximal and distal debris flows), and background sedimentation of fines by settling from suspension. The resulting association, called here early TST, is made up of three main facies stacked in a retrogradational pattern: (i) clast-poor, massive to poorly stratified diamictites; (ii) thinly bedded diamictites; and (iii) shales with IRD. The upper part of this type of TSTs (complete TSTs) is made up of open-marine, IRD-free shales, unusually associated with fine-grained sandstones with bi- or unidirectional ripples indicative of tenuous wave reworking in shallow-marine environments. Clastic dilution owing to high sediment contributions would prevent accumulation of significant volumes of organic matter in the shales. 2. Base-cut TSTs: Where the basal transgressive portion is mostly omitted, the early TST is represented, from base to top, by distinctive distal debris-flow deposits (in some cases associated with low-density Bouma-type turbidites) and shales with IRD. The upper part of the TST is represented by basinal (below wave base) IRD-free shales. Alternatively, in extreme settings with a very high sea-level rise rate, the entire TST is made up exclusively of IRD-free shales (the basal shale model of Wignall,

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Figure 16. Postglacial transgressive systems tract (TST) spectrum, based on relative influence of sediment supply and sea-level-rise rates. See text for further discussion. Grain-size scale in logs: Si—silt; S— sand; G—gravel. LST—lowstand systems tract.

1994). Wave reworking is unlikely in these fast transgressions. These base-cut TSTs are common in many glacial-postglacial transitions and basically reflect the drastic sea-level rise related to ice melting. Owing to a relatively low sediment supply rate with respect to the sea-level-rise rate, organic-rich shales could be deposited and preserved. Water Salinity during Deglaciation Widespread evidence indicates that fresh-water contributions into the inland seas developed in early postglacial times. This fresh-water incursion, as melting icebergs, is interpreted to have had the combined effect of creating a layer of brackish water and a relatively high suspended-sediment load, as described for fjords in Greenland by Syvitski et al. (1987, 1996). Three independent lines of evidence suggest drastic fluctuations of salinity from normal marine through brackish to fresh waters for these postglacial seas: (1) ichnofacies associations, (2) associated fauna, and (3) geochemical characteristics. An inland sea paleogeographic model has been proposed for the Transantarctic Basin (Barrett et al., 1986; Collinson and Miller, 1991; Miller and Collinson, 1994; Collinson et al., 1994). Early Permian palynoflora (fossil spores and pollen), recorrelated

with Australian palynomorph zones, is present in sediments resting on glaciogenic diamictites (Pagoda Formation). This palynoflora indicates that the glaciation in the Transantarctic Mountains was restricted to the Asselian–Sakmarian (Isbell et al., 2005). These deposits above the glaciogenic diamictites are grouped under the Mackellar Formation of the Central Transantarctic Mountains (Fig. 1). This unit is mostly made up of black, finely laminated shales and subordinate fine-grained sandstones with ripples, resting on the glaciogenic deposits of the Pagoda Formation (Lindsay, 1970; Miller, 1989); the upper Pagoda Formation is characterized by massive shales with IRD (Miller et al., 1991). C/S ratios in these sediments are extremely high, indicative of fresh to slightly brackish waters (Miller et al., 1991); TOC content is low (<1%, exceptionally ~2%) and is dominated by terrestrial plants (type III). This last example, in addition to those documented herein in more detail, indicates that early postglacial times were characterized by inland seas with dominantly brackish waters rather than lakes. Alternative views sustain the presence of huge lakes rather than inland seaways. This presumption is not based on detailed paleogeographic studies that demonstrate that closed bodies of water (lakes) without connection to the open sea prevailed during postglacial times but rather on indirect evidence of brackish to fresh-water salinities provided by geochemical studies and, consequently, to the extreme scarcity to absence of marine fossils. A good example of that debate is the Karoo Basin during the glacialpostglacial transition. Owing to the presence of marine fossils at the top of the Dwyka Formation (McLachlan and Anderson, 1973) along the western and northern margins of the basin, there seems to be agreement for the existence of a marine transgression related to eustatic sea-level rise following rapid global deglaciation. However, nonmarine conditions are suggested for the overlying Prince Albert, Whitehill, and Collingham Formations by Milani and De Wit (2008), who further speculate on nonmarine conditions for the Dwyka diamictites by stating that the “water-lain diamictites, previously modeled as deposited in a wide marine-shelf environment (Visser 1989, 1997), may represent glaciogenic lake sediments deposited beneath and peripheral to the major continental ice sheet that covered much of Gondwana at that time” (p. 333). In this interpretation the postglacial marine mudstones were part of “a short marine transgression related to eustatic sealevel rise following rapid global de-glaciation” (p. 333). Fluctuating salinity (from brackish to fresh-water conditions) actually prevailed during early postglacial times, as evidenced mostly by geochemical data, and similar fluctuating salinity conditions probably persisted throughout the sedimentation of the Prince Albert, Whitehill, and Collingham Formations. The geochemical data supporting events of algal blooms in brackish to fresh water for the Whitehill Formation (Faure and Cole, 1999) do not need to be necessarily related to a landlocked inland body of standing water, the definition of a lake (cf. Bates and Jackson, 1984), but rather to an inland sea whose salinity may have been still controlled by fresh-water incursions related to retreating valley glaciers along the northen margin of the Karoo Basin. The retreat of

Transgressions related to the demise of the Late Paleozoic Ice Age marine conditions in the Karoo Basin was closely related to the emerging Cape fold belt along its southern margin by Middle to Late Permian time, when deep-water deposits were replaced by deltaic sediments, followed by fluvial and lacustrine sediments (Wickens, 1992; Veevers et al., 1994). Cataclysmic floods (jökulhlaups) related to the Pleistocene deglaciation also have been invoked to explain the widespread presence of areally extended fine-grained sediments deposited in fresh waters. After an ice sheet retreated in a terrestrial setting, meltwater was ponded in proglacial lakes developed behind terminal moraines. As lake level rose with deglaciation, breaching of the moraines occurred, flooding vast areas. This process has been identified in several areas that were affected by the Pleistocene glaciation (for a review, see Baker, 1997, and references therein). However, this model, when applied to many of the glacial-postglacial sections in the basins affected by the LPIA, fails to explain the scarcity of facies that could be assigned to marginal lacustrine settings. On the contrary, open-marine fossiliferous shales present in the postglacial TSTs analyzed in this contribution seem to be interbedded with the allegedly lacustrine, deep-water (below wave base) deposits without any intervening shallower water facies (i.e., Malanzán Formation in the Paganzo Basin, Lower Barakar Formation in the Satpura Basin). Postglacial Transgressions and Source Rocks Early postglacial sea-level rise favored creation of accommodation space with preservation potential for productivityanoxia events and peat-forming conditions by rapid water-table rise. Total organic content is a function of the rate of diluting sedimentation, the rate of degradation, and the onset of low-oxygen, anoxic conditions (Wignall, 1994). These conditions, particularly low sediment input and anoxic bottom conditions, are produced by the increasing water depth during a transgression. Therefore, the postglacial TST includes the condensed section that represents conditions of offshore sediment starvation and consequently optimal conditions for the accumulation and preservation of types I and II organic matter (Loutit et al., 1988; Schlutter, 1998). Finegrained facies in a TST contain high values of TOC and high HI (hydrogen index) values and may be more prone to anoxia (Wignall, 1991; Pasley et al., 1991, 1993). Conditions of clastic starvation are more likely during rapid transgressions triggered by high rates of relative sea-level rise, such as in the postglacial transgressions analyzed herein. The shales in most postglacial TSTs seem to fit the model of condensed-section shales, deposited during a period of the most rapid rise in sea level when clastic contributions are trapped updip. A similar context has been proposed for the earliest Silurian (Rhuddanian) postglacial hot shales of North Africa (Lüning et al., 2000). The uneven lateral extent of these deglacial, transgressive black shales has been linked to their deposition within a glacially sculpted topography. During the early stages of this transgression, triggered by a rapid ice sheet recession, the most deeply scoured areas of the postglacial shelf were the first to be flooded. Black shales were

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then deposited in the paleovalleys, which became isolated basins where conditions for the generation of low-oxygen conditions were more favorable. Additionally, some paleolatitudinal constraints seem to have been common for the presence of tasmanite beds. As pointed out by Revill et al. (1994), sediments where Tasmanites is abundant (tasmanite beds), such as in the Late Ordovician– Early Silurian of North Africa, the Devonian of Brazil, and the Late Carboniferous–Early Permian tasmanites of Tasmania, seem to have had similar high (~70°–75°) paleolatitudes, inferred from paleomagnetic work (Smith et al., 1981; Bachtadse and Briden, 1990; Li and Powell, 2001) and continental reconstructions based on climatically sensitive lithofacies (Scotese and Barrett, 1990). Evolution of Basin Margins after Postglacial Transgressions In general, the postglacial TSTs across Gondwana are succeeded by progradational deposits (HSTs) as the rate of sea-level rise and accommodation space decreased (Fig. 2). Facies variability is the norm within this common progradational stacking pattern, ranging from forced regressions (sensu Posamentier et al., 1992) dominated by the drastic intrusion of coarse clastics (Paganzo Basin, western Argentina, Limarino et al., 2010) to gradual filling of the basin by prodelta to delta-front fine- to medium-grained sandstones and mudstones (Calingasta-Uspallata Basin, western Argentina), in some cases with significant tidal influence and associated delta plain coals (Barakar Formation in India). This wide gamut of facies in the overlying progradational association contrasts sharply with the narrower spectrum of facies in the underlying postglacial TST, which seems to be the product of a relatively simpler interaction between the rate of sediment supply and the rate of sea-level rise associated with glacier retreat. In exceptional cases, such as along the southern margin of the Karoo Basin, where tectonic subsidence rates outstripped sediment supply rates along a foredeep next to an emerging mountain range (Cape fold belt), the deepening of the basin continued, and the postglacial TST was overlain by basinal organic-rich shales (Whitehill Formation) and deep-water, low- and high-density turbidites (Collingham, Laingsburg, and Ripon Formations) (Catuneanu et al., 2002). In this extreme context the rate of sediment supply was unable to keep up with the rate of new accommodation space created by a very high subsidence rate, creating conditions of basin underfilling (forced transgression of Chough and Hwang, 1997) after the effects of glacioeustatic sea-level rise ceased (cf. Figs. 2C and 2D). CONCLUDING REMARKS The examples of glacial-postglacial transitions reviewed herein cover a wide range of tectonic settings and basin types, from backarc foreland basins to rifts. In all of them a clear retrogradational stacking pattern is detectable in the transition from

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glacially dominated settings to glacially influenced early postglacial environments. Drastic landward facies shifts are indicative of considerable relative sea-level rises, whose main driving mechanism in these cases is glacioeustasy. Although tectonically induced subsidence has to be considered an additional element in the relative sea-level rise inferred from these sections, these postglacial transgressions cannot be explained primarily by the sole effect of thermal sag-like or tectonic subsidence and consequent flooding. The resulting stratigraphic record can be framed in sequence stratigraphic terms as postglacial transgressive systems tracts (TSTs) principally driven by glacioeustasy. The balance between sediment supply and sea level is critical to the deposition of organic-rich black shales in the postglacial TSTs. A relatively high sediment supply rate can produce significant clastic dilution and prevent starved conditions in the newly inundated postglacial shelf; on the other hand, this negative influence could be counterbalanced by the rapid eustatic sea-level rise produced during deglaciation. The two types of TSTs defined in this contribution (complete TST and base-cut TST) can be considered end members, reflecting these two possible scenarios. ACKNOWLEDGMENTS This work benefited greatly from helpful discussions through the years with many colleagues, particularly Arturo Amos, Jim Collinson, John Crowell, and Johan Visser. The author thanks Antonio Rocha-Campos and John Veevers for their insightful reviews of the manuscript. REFERENCES CITED Adie, R.J., 1952, The position of the Falkland Islands in a reconstruction of Gondwanaland: Geological Magazine, v. 89, p. 401–410, doi: 10.1017/ S0016756800068102. Ahmad, N., 1975, Son Valley Talchir glacial deposits, Madhya Pradesh, India: Journal of the Geological Society of India, v. 16, p. 475–484. Aitken, J.F., Clark, N.D., Osterloff, P.L., Penney, R.A., and Mohiuddin, U., 2004, Regional core-based sedimentological review of the glaciallyinfluenced Permo-Carboniferous Al Khlata Formation, South Oman Salt basin, Oman: GeoArabia, v. 9, p. 16. Aksu, A.E., 1984, Subaqueous debris flow deposits in Baffin Bay: Geo-Marine Letters, v. 4, p. 83–90, doi: 10.1007/BF02277077. Al-Belushi, J., Glennie, K.W., and Williams, B.P.J., 1996, Permo-Carboniferous glaciogenic Al Khlata Formation, Oman: A new hypothesis for origin of its glaciation: GeoArabia, v. 1, p. 389–403. Amos, A.J., 1980, La fauna de invertebrados en la cronología del Carbónico y Pérmico de Argentina: Buenos Aires, Congreso Argentino de Paleontología y Bioestratigrafía, 2nd, and Congreso Latinoamericano de Paleontología, 1st, Actas 4, p. 231–234. Amos, A.J., and López-Gamundí, O.R., 1981a, Late Paleozoic tillites and diamictites of the Calingasta-Uspallata and Paganzo basins, in Hambrey, M., and Harland, W., eds., Earth’s Pre-Pleistocene Glacial Record: Cambridge, UK, Cambridge University Press, p. 859–868. Amos, A.J., and López-Gamundí, O.R., 1981b, Late Paleozoic Sauce Grande Formation of eastern Argentina, in Hambrey, M., and Harland, W., eds., Earth’s Pre-Pleistocene Glacial Record: Cambridge, UK, Cambridge University Press, p. 872–877. Amos, A.J., and Rolleri, E., 1965, El Carbónico marino en el valle CalingastaUspallata, San Juan-Mendoza: Boletín de Informaciones Petroleras, v. 368, p. 50–71. Andreis, R.R., 1984, Análisis litofacial de la Formación Sauce Grande (Carbónico superior?) Sierras Australes, provincia de Buenos Aires: San Car-

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MANUSCRIPT ACCEPTED BY THE SOCIETY 21 DECEMBER 2009

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The Geological Society of America Special Paper 468 2010

From bergs to ergs: The late Paleozoic Gondwanan glaciation and its aftermath in Saudi Arabia John Melvin Ronald A. Sprague Christian J. Heine Saudi Aramco, Box 12646, Dhahran 31311, Saudi Arabia

ABSTRACT The late Paleozoic (Carboniferous–Permian) Gondwanan glaciation is represented in the subsurface of eastern and central Saudi Arabia by the Hercynian (or pre-Unayzah) unconformity and the lower part of the overlying Unayzah Formation. The subsequent postglacial transgression is manifest in the upper members of the Unayzah Formation as well as the lowermost clastic deposits of the overlying Khuff Formation. Its component sediments result from ongoing climatic amelioration following the demise of the ice age, as well as tectonic influences related to the creation of the Neotethys Ocean. The Unayzah Formation is subdivided into four stratigraphic members. Thus, directly overlying the Hercynian unconformity in many places are sandstones and minor conglomerates of the Unayzah C member. These were laid down within a widespread, braided glaciofluvial depositional system. They represent glacial outwash produced during times of glacial retreat throughout the duration of the late Paleozoic glaciation. An unknown number of glacial readvances occurred that significantly deformed these retreat-phase outwash sands and gravels, creating major glacially tectonized push moraine nappes. Those are interpreted from a number of distinctive and discrete shear zones that are uniquely associated with the Unayzah C member. The upper surface of the Unayzah C member is an unconformity that marks the final subglacial surface at the time of maximum advance of the ice. The terminal melt-out phase of the Gondwanan glaciation is represented by the Unayzah B member. Paleomagnetic evidence suggests that this member was deposited at high latitudes, ~75° S. This member comprises a large number of depositional facies that are essentially glaciolacustrine in character. Those facies include (1) smallscale ice-contact push moraines indicative of minor glacial readvance, (2) ice-proximal sublacustrine debris flows (massive diamictites) and associated gravity flow deposits, and (3) ice-distal, sublacustrine stratified diamictites, ripple cross-laminated sandstones, and laminated mudrocks. Facies associations within the Unayzah B member consistently show evidence of sustained glacial retreat and flooding of the landscape by filling and spilling over of numerous glacial lakes. This flooding sequence probably

Melvin, J., Sprague, R.A., and Heine, C.J., 2010, From bergs to ergs: The late Paleozoic Gondwanan glaciation and its aftermath in Saudi Arabia, in LópezGamundí, O.R., and Buatois, L.A., eds., Late Paleozoic Glacial Events and Postglacial Transgressions in Gondwana: Geological Society of America Special Paper 468, p. 37–80, doi: 10.1130/2010.2468(02). For permission to copy, contact [email protected]. ©2010 The Geological Society of America. All rights reserved.

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Melvin et al. represents the maximum climatically related postglacial transgressive event in Saudi Arabia. In the western part of the study area there is evidence that the ice remained longer, and it is tentatively interpreted as a local center for high altitude (“alpine”) glaciation. Deposition of the Unayzah B member was terminated abruptly by a drainage event that is marked by a widespread sharp contact with the overlying unnamed middle Unayzah member. The latter member displays no unequivocally glacially related depositional facies, and paleomagnetic data suggest that it was deposited at ~55° S. It is dominated by red floodplain siltstones and very fine–grained sandstones that contain relatively isolated bodies of coarser fluvial and eolian sandstones. These eolianites display possible cold-climate characteristics. This is particularly true in the western part of the study area, which is consistent with a relatively sustained, high altitude ice cap in that area. The unnamed middle Unayzah member is capped in many places by a paleosol horizon. This represents a hiatus of unknown but probably prolonged duration and thus suggests a disconformable contact between the unnamed middle Unayzah member and the overlying Unayzah A member. The Unayzah A member is dominated by sediments that are strongly characteristic of terrestrial deposition in a semiarid to arid environment (including ephemeral lakes and streams as well as eolian deposits). Paleomagnetic data suggest paleolatitudes ~28° S. The continental eolian clastic deposits of this member in places display a cyclicity in their stratal architecture that is related to fluctuations in the paleo–water table. These fluctuations are possibly related to distant marine transgression, which is supported by the occurrence of a distinctive bioturbated sandstone very close to the top of the Unayzah A member. That marine-influenced sandstone is observed in widely separated localities at either end of the study area and may represent the final breakthrough of transgressive marine waters close to the end of Unayzah A time. In several places the uppermost deposits of the Unayzah A member are characterized by thick paleosols. These represent a prolonged period of nondeposition, interpreted to be directly related to thermal doming of the Arabian plate prior to rifting and opening of the Neotethys Ocean, and the consequent formation of the preKhuff unconformity, which terminated Unayzah deposition. Overlying the pre-Khuff unconformity are various siliciclastic facies of the eponymous Basal Khuff Clastics member of the Khuff Formation. The depositional facies of the lowermost Basal Khuff Clastics range from shallow marine in the southeastern part of the study area to predominantly fluvial in the west. This reflects the westward-directed transgression of the Khuff Formation following thermal collapse in the wake of the rifting that created the Neotethys Ocean. That tectonically related transgression reached its fullest expression with deposition of the carbonates and evaporites that dominate the upper members of the Khuff Formation. This stratigraphic evolution of the late Paleozoic in Saudi Arabia can confidently be correlated in sequence stratigraphic terms with coeval sediments laid down across the Arabian Peninsula.

INTRODUCTION The Arabian plate was subjected to a major episode of tectonic deformation during the middle Carboniferous (Husseini, 1992; McGillivray and Husseini, 1992; Wender et al., 1998; Al-Hajri and Owens, 2000; Konert et al., 2001; Sharland et al., 2001; Al-Husseini, 2004). There followed a period of erosion that lasted from Namurian to early Westphalian (Serpukhovian to Bashkirian) times (Al-Husseini, 2004), and which is represented in the stratigraphic record by the Hercynian (or pre-Unayzah)

unconformity in Saudi Arabia. This unconformity is widely recognized in the subsurface across the study area of eastern central Saudi Arabia (Fig. 1A), and it truncates older formations ranging in age from Proterozoic to Early Carboniferous. This tectonic event was closely followed by the inception of the Late Paleozoic Ice Age in Arabia (Sharland et al., 2001; Al-Husseini, 2004). The depositional products of that late Paleozoic Gondwanan glacial episode were laid down in the first instance upon the Hercynian unconformity. The subsequent postglacial transgression was affected during the Permian by tectonic events that related to

Late Paleozoic Gondwanan glaciation in Saudi Arabia

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Figure 1. Location maps. (A) Map of eastern central Saudi Arabia, showing the locations of the wells studied. Note also the location of the Al Batin Arch (Faqira et al., 2009) in the western part of the study area. (B) Map of the Arabian Peninsula, showing the principal study area for this paper, as well as areas in Oman (A), Yemen (B), and southwestern Saudi Arabia (C), where glaciogenic sediments of Carboniferous–Permian age crop out.

opening of the Neotethys Ocean. The geological record of these events is preserved in Saudi Arabia in the siliciclastic deposits of the Unayzah Formation and the clastic rocks that occur at the base of the overlying Khuff Formation in Saudi Arabia, and culminates in the carbonates that dominate in the higher parts of the Khuff Formation. The depositional record of the Gondwanan glaciation on the Arabian plate has been well documented from Oman (Braakman et al., 1982; Hughes Clarke, 1988; Levell et al., 1988; Al-Belushi et al., 1996; Aitken et al., 2004; Osterloff et al., 2004a) as well as from northern Yemen (Roland, 1979; Kruck and Thiele, 1983) (Fig. 1B). In Saudi Arabia, boulder-bearing diamictites of inferred glacial origin have been recognized in the Juwayl Formation

where it crops out in the southwestern Wajid region (Helal, 1964; Kellogg et al., 1986; McClure, 1980; McClure and Young, 1981; McClure et al., 1988) (Fig. 1B) and dated as Late Carboniferous (Stephanian = Moscovian to Gzhelian) to Early Permian (Sakmarian) in age (Cameron and Hemer in McClure, 1980; McClure and Young, 1981). Since the late 1980s a large number of wells have penetrated the Khuff and Unayzah Formations in the subsurface of Saudi Arabia as part of an aggressive exploration and exploitation program in the search for oil and non-associated gas. The recovery of many thousands of feet of core in that drilling effort has enabled much progress to be made in understanding the origin of these formations and their constituent members, as well as their stratigraphic distribution (architecture) in the subsurface

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of central and eastern Saudi Arabia. Detailed studies are ongoing, and some results have already been released (e.g., Melvin and Heine, 2004; Melvin et al., 2005; Melvin and Sprague, 2006). The current paper discusses the nature, origin, and stratigraphy of siliciclastic deposits representing the Unayzah Formation in the subsurface of east-central Saudi Arabia, as well as overlying siliciclastic sediments that occur at the base of the otherwise predominantly carbonate Khuff Formation. The rocks are described and considered in terms of their relationship to the Gondwanan glaciation per se, and also, as appropriate, in terms of the extent to which they reflect the development of the late Paleozoic, postglacial history of siliciclastic sedimentation in Saudi Arabia. Twenty-four representative cored wells were investigated (Fig. 1A), and a total of 6351 ft (1905.3 m) of core was described and interpreted. The cored intervals were integrated with conventional wireline logs, and a stratigraphic framework has been constructed that places the various core-derived depositional facies in their architectural context. STRATIGRAPHIC SETTING Sharland et al. (2001) presented an evaluation of the complete sequence stratigraphy of the entire Arabian plate. Those authors recognized a number of “Tectonostratigraphic Megasequences” (TMS) that provided the foundations for that sequence stratigraphic framework. Specifically, their TMS AP5 megasequence spans the period from the middle Carboniferous “Hercynian” tectonic event to the early Late Permian rifting event that created the Neotethys Ocean (cf. Bishop, 1995; Loosveld et al., 1996) (Sharland et al., 2001). TMS AP5 is the last TMS to be dominated by siliciclastic sediments prior to the movement of the plate into subtropical latitudes where carbonate and evaporite deposition became dominant (cf. Beydoun, 1991; Al-Fares et al., 1998). The base of this megasequence is marked by the Hercynian unconformity (Fig. 2). The top of TMS AP5 is the pre-Khuff unconformity of Senalp and Al-Duaiji (1995), Evans et al. (1997), and Wender et al. (1998) (Sharland et al., 2001). The current paper will therefore discuss in its entirety the TMS AP5 megasequence of Sharland et al. (2001) (viz. the Unayzah Formation), as well as the lowermost siliciclastic deposits of the superseding TMS AP6 megasequence, namely the Basal Khuff Clastics member of the Khuff Formation that abruptly overlies the pre-Khuff unconformity. Lithostratigraphy The Unayzah Formation in the subsurface of eastern central Saudi Arabia sits directly upon the Hercynian unconformity. Historically, it has been subdivided informally into three members, the Unayzah A (youngest), B, and C (oldest) (Ferguson and Chambers, 1991; McGillivray and Husseini, 1992). It varies considerably in thickness, with the greatest thickness variations taken up in the Unayzah B and C members. The Unayzah A member is truncated by the pre-Khuff unconformity (Senalp and

Al-Duaiji, 1995). This is recognizable throughout the subsurface (Al-Husseini, 2004) and is overlain by interbedded sandstones, shales, and thin carbonates of the Basal Khuff Clastics member of the Khuff Formation (Fig. 2). Notwithstanding the tripartite subdivision of the subsurface Unayzah by Ferguson and Chambers (1991), until very recently there had existed no rigorous definition of its constituent members. This lack of definition arises from the intrinsic variability in the wireline-log character of the stratigraphic units among the wells. That variability stems from extreme variations in thickness and facies among the members of the Unayzah, as well as from an abundance of erosional and nondepositional hiatuses throughout the formation. Melvin and Sprague (2006), in an extensive corebased study, revisited the lower Unayzah (i.e., the pre-Unayzah A part of the formation). Following a detailed description of the various depositional facies within these rocks, they redefined the stratigraphy with the introduction of a new, “un-named middle Unayzah member” that occurs between the Unayzah B member and the Unayzah A member of Ferguson and Chambers (1991) (Fig. 2). The continued extensive use of cores and core-related information from the present study has enabled further consolidation of our understanding of the lithostratigraphic relationships within and among the members of the Unayzah and the overlying Basal Khuff Clastics. All of these members of the Unayzah Formation in the subsurface of Saudi Arabia continue to be represented by an informal working nomenclature. Given the enhanced understanding of these rocks that is in large part represented herein, the need for a revised, formalized lithostratigraphic nomenclature scheme within the Unayzah has been recognized and is being addressed by Saudi Aramco. The results of that reexamination of the stratigraphic nomenclature will be published in a separate, future paper. Biostratigraphy It is extremely difficult to obtain definitive age dates from the Carboniferous-Permian succession of the Arabian plate, and in particular in Saudi Arabia from the Unayzah Formation. There, the primary biostratigraphic tool has been limited to palynology, owing to a lack of macrofossils in these predominantly terrestrial rocks. Furthermore, many well sections within the Unayzah are barren of any fossils, including palynomorphs. Nonetheless, following earlier work by Stephenson (1998), Al-Hajri and Owens (2000), and Stephenson and Filatoff (2000a, 2000b), Stephenson et al. (2003) produced a palynological zonation scheme that enables correlation to a varying degree among all the various members of the Unayzah Formation, as well as the Basal Khuff Clastics, in Saudi Arabia, with the Al Khlata, Gharif, and the lower part of the Khuff Formations of Oman (Fig. 2). This palynozonation is supported in part by paleontological studies carried out in Oman by Angiolini et al. (2006) and was recently updated and modified by Stephenson (2006). Thus, the Unayzah C has been assigned to the OSPZ1 Zone by Stephenson et al. (2003), and is considered by those authors to be Late Carboniferous

Late Paleozoic Gondwanan glaciation in Saudi Arabia

A

B

Changhsingian Wuchiapingian

KSA subsurface stratigraphy TMS* (Sharland et al., 2001)

TMS AP6

This paper

Middle

Wordian Roadian

hiatus

Kungurian

Khuff

hiatus

Early

Unayzah A member

Oman lithostratigraphy Levell et al. (1988)

Osterloff et al. (2004a, 2004b)

Lower Khuff OSPZ 6

OSPZ 5

OSPZ 4

DS P17 thru DS P19

Upper DS P15 Gharif

Middle DS P13 Gharif

hiatus

Artinskian

hiatus

Unnamed middle Unayzah member

Sakmarian Asselian

Stephenson et al. (2003) and Angiolini et al. (2006)

Khuff Basal

Capitanian

KSA / Oman biostratigraphy

Gharif Formation

Late

Stratigraphic units

PERMIAN

41

TMS AP5

Unayzah B member

OSPZ 3

c b a Rahab Shale

OSPZ 2

Extensive glacio-lacustrine

Lwr Gharif DS P10 DS P8 DS P6

DS CP Gzhelian

Unayzah C member (multiple hiatuses)

Moscovian

Late

CARBONIFEROUS

hiatus

OSPZ 1 Lower Al Khlata

DS C30 Hercynian hiatus

Unconformity hiatus

Figure 2. Late Carboniferous-Permian stratigraphy of Saudi Arabia (KSA) and Oman. Note: (A) Tectonostratigraphic Megasequences (TMS) of Sharland et al. (2001); (B) biostratigraphic zonation (OSPZ1–OSPZ6) of Stephenson et al. (2003) that provides critical linkage between the successions of Saudi Arabia and Oman.

(Stephanian = Moscovian to Gzhelian) in age. It is equated with the lower part of the Al Khlata Formation in Oman, specifically depositional sequence DS C30 and the lower part of DS CP of Osterloff et al. (2004a) (Fig. 2). Stephenson and Filatoff (2000a), Stephenson et al. (2003), and Stephenson (2004) showed that the Unayzah B member is Early Permian (Asselian to Sakmarian) in age and is characterized by palynomorph assemblages of the OSPZ2 Zone. This is equated with the upper part of the Al Khlata Formation of Oman (including the Rahab Shale of Levell et al.,

1988). Those rocks are defined by Osterloff et al. (2004a) as the upper part of depositional sequence DS CP, as well as DS P6 and DS P8 in Oman (Fig. 2). The correlation of rock units between Saudi Arabia and Oman using the OSPZ3 Zone (and its three subbiozones) is more challenging. These subbiozones are restricted in occurrence, and so may not be recognizable throughout Arabia owing to either paleophytogeographical variation or hiatus (Stephenson et al., 2003). Thus, although the Haushi Limestone at the

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Melvin et al.

top of the lower Gharif member in Oman was shown by Angiolini et al. (2006) to correlate with the OSPZ3c Subbiozone, in Saudi Arabia Stephenson et al. (2003) considered the OSPZ3 Biozone to be absent, at least in part, and so to represent a depositional hiatus at the time of deposition of the lower Gharif member. Melvin and Sprague (2006) suggested that their “unnamed middle Unayzah member” of the Unayzah in Saudi Arabia may be at least partially the terrestrial lateral equivalent of the marine lower Gharif member of Oman (Fig. 2). This would give it some degree of equivalence with the OSPZ3 Biozone of Stephenson et al. (2003) and suggests correlation in Oman with depositional sequence DS P10 of Osterloff et al. (2004b). In this context, Stephenson (2006, written commun.) acknowledged that in Saudi Arabia the OSPZ3 Zone, rather than being absent, could alternatively be placed “against some barren section” in the Unayzah. Any hiatus in Saudi Arabia thus would have been of more limited duration than originally implied by Stephenson et al. (2003). Angiolini et al. (2006) demonstrated that the lower Gharif member in Oman is late Sakmarian in age; by implication this age would also apply to the unnamed middle Unayzah member. The Unayzah A member in Saudi Arabia has been broadly correlated with the OSPZ4 Biozone (and possibly the OSPZ3b and OSPZ3c Subbiozones) by Stephenson et al. (2003). This correlation suggests a general equivalence to the middle Gharif member of Oman (depositional sequence DS P13 of Osterloff et al., 2004b) (Fig. 2). Stephenson et al. (2003, p. 486) observed that the palynological changeover between OSPZ4 and OSPZ5 is “probably the greatest recorded in the Permian palynological succession in Oman and Saudi Arabia.” They also noted how in Oman there is an abrupt change in depositional environment, as well as in chemostratigraphy and heavy mineral assemblages, between the middle and upper members of the Gharif Formation. Furthermore, there is evidence for local incision and paleosol development in the upper part of the middle Gharif member (Stephenson et al., 2003; Osterloff et al., 2004b), which suggests a significant depositional hiatus between the middle and upper Gharif members. It will be demonstrated in the present paper that similar changes occur between the top of the Unayzah A member and the Basal Khuff Clastics, and that they are thus similarly separated by an unconformity of considerable significance. The Basal Khuff Clastics and the upper Gharif member are therefore probably equivalent lithostratigraphic units, as was implicit in stratigraphic diagrams presented by Sharland et al. (2001, p. 85) and Stephenson et al. (2003, p. 470). These units are not, however, time-equivalent, as the upper Gharif member is characterized by the OSPZ5 Palynozone, whereas the Basal Khuff Clastics member is definitively represented by the younger OSPZ6 Zone (Stephenson, 2006). This diachroneity is considered by the present authors to be a reflection of the onset in Oman of the major transgression that culminated in widespread deposition of carbonates of the Khuff Formation and which is related to the opening of the Neotethys Ocean.

UNAYZAH C MEMBER Depositional Characteristics A total of 1160 ft (348 m) of core from the Unayzah C member was examined in detail in this study. Characteristically, it is a hard, quartz-cemented sandstone unit that is extremely variable in thickness, reflecting deposition upon a strongly undulating surface (the Hercynian unconformity). Although absent in some places, it can be present as several hundred feet of sediment only a few kilometers away. In general, it displays the most distinctive and reproducible wireline-log signature of any of the members of the Unayzah Formation in the Saudi Arabian subsurface. Thus it is characterized by a monotonous, very low API gamma-ray log from most wells, and similarly an essentially featureless sonic log with very low (fast) sonic transit times. These featureless wireline traces are interrupted locally by high gamma and slow sonic spikes, respectively. In core, the sandstones are arranged in multistory bedsets that are commonly several tens of feet thick (Fig. 3). Within these bedsets, individual beds pass upward from sharp, erosional basal surfaces, locally strewn with intraformational mudclasts and pebbles and cobbles of both massive and laminated sandstone (Fig. 3A, 10.5–12.5 ft; Fig. 3B, 1–9 ft; Fig. 4A), into upward-fining, poorly to moderately sorted, very coarse– to medium-grained sandstones. Primary depositional structures are hard to identify in these sandstones, although in places the rocks may display low-angle planar lamination, or less commonly trough cross-lamination (Fig. 3). In some wells (e.g., well 13) beds within the Unayzah C member contain abundant mud clasts, up to 3 cm long, that are uniformly dispersed throughout the bed, and oriented more or less parallel to bedding (Fig. 3B). Notwithstanding the above-described features, very commonly the rock is either “massive” or characterized only by an extremely diffuse, poorly sorted fabric that passes up into discontinuous, crinkly, argillaceous laminations that may be enhanced by stylolitization (see below). Rarely, thin (centimeter scale) beds of gray-green, sandy siltstone separate the sandstone bedsets. The uppermost contact of the Unayzah C member has been cored in a number of wells, and it is in all cases extremely sharp. Melvin and Sprague (2006) described how in well 8, the contact is also extremely irregular, with reentrants that are infilled by sediment of the overlying Unayzah B member. Beneath that irregular contact, the uppermost 1 ft (30 cm) of the Unayzah C member in this well displays a high degree of brecciation, with a network of abundant, intersecting clay-lined fractures. This has been interpreted to be an immature paleosol horizon (Melvin and Sprague, 2006). In most wells where the upper contact has been cored, the Unayzah C member is overlain by the Unayzah B member (e.g., Figs. 3A and 3B). In well 11 it is directly overlain by distinctive facies of the Unayzah A member (Fig. 3C). All of these stratigraphic relationships at the top of the Unayzah C member point strongly to that contact as an unconformity, the significance of which is discussed below.

Figure 3. Core logs from wells 7, 13, and 11, showing the fundamental character of the Unayzah C member. Note that in all wells, the abundance of “structural features” (stylolites, fractures, etc.) is highlighted in gray, and that these are uniquely limited to the Unayzah C member. In addition, note the intensely sheared zones in well 7 (between 0 and 10 ft) and well 13 (19–20 ft, 57–60 ft, 75–85 ft). The rocks in well 13 are characterized by an abundance of mud clasts dispersed throughout the beds. The Unayzah C member can be overlain by sediments of either the Unayzah B member (wells 7 and 13) or the Unayzah A member (well 11). The key to all symbols used in core logs in this figure and all other relevant figures throughout the paper is presented in Appendix Figure A1.

Figure 4. Characteristics of glaciotectonic shear zones within the Unayzah C member. (A) Core photograph illustrating a typical shear zone within the Unayzah C member in well 7 (cf. Figure 3A, 0–10 ft). This occurs between 14,108 ft and 14,115 ft, and displays evidence for both brittle and ductile shear. It is overlain by intraformational conglomerate of mud clasts and laminated sandstones (14,104.5 ft to 14,106.5 ft), followed by undeformed massive sandstones from 14,104.5 ft to the top of the core. (B) Gamma-ray log through the Unayzah C member in well 14, showing how spiky, high API values are associated with shear deformation in core. The core examples B(i) and B(ii) are arrowed on the gamma-ray log. The otherwise monotonous, low API gamma trace reflects the relatively featureless rock that characterizes the Unayzah C member away from the shear zones (example at B[iii] is also arrowed on the wireline log).

Late Paleozoic Gondwanan glaciation in Saudi Arabia Post-Depositional Characteristics In many of the wells that have penetrated and cored the Unayzah C member (e.g., wells 6, 7, 8, 11, 13, 14, 17, and 24), distinctive zones of intense subhorizontal shear have been observed (Melvin and Sprague, 2006). These shear zones occur with associated low-angle overfolds and range in thickness from 2 ft (0.6 m) to >20 ft (6 m). Examples are illustrated in Figures 3A, 3B, and 4. The shear zones display a complex variety of features ranging from brittle shear to ductile shear to simple softsediment deformation. These zones are characteristically overlain and underlain by much thicker intervals of sandstone that are devoid of any such deformation (Figs. 3A and 4B). The shear zones are highly distinctive wherever they are seen in core, and all these zones recorded to date clearly correlate with intervals that display a spiky, high gamma profile on wireline logs (e.g., Fig. 4B). Commonly, they are also associated with a strongly chaotic signature on image logs (M.H. Prudden, 2005, personal commun.). This has enabled tentative identification of the shear zones in uncored wells: in such cases the chaotic zones are separated by considerable thicknesses (several tens of feet) of apparently normally stratified sediment, as has been observed in the cored examples. Origin and Evolution of the Unayzah C Member The character of the multistory bedsets of the Unayzah C member suggests that it was deposited in an environment that was dominated by processes of repeated erosion and deposition. The multiple scoured contacts imply a strongly channelized environment, and the grain size and primary sedimentary structures are suggestive of high energy conditions, as indeed are the thick beds with abundant dispersed mud clasts. The diffuse fabric seen in many of these sandstones is ascribed to dewatering, and the associated argillaceous crinkly laminations probably represent elutriated fines associated with that dewatering. The depositional system of the Unayzah C member was thus dominated by an abundance of channels characterized by rapid, high energy, sandy bedload sedimentation—i.e., an extensive alluvial braided plain. The thin siltstone beds represent rare low-stage suspension fallout deposits of fine-grained sediment upon the upper surfaces of the channel braided bars. As has been discussed above, limited biostratigraphic evidence suggests that the Unayzah C member is Late Carboniferous (Stephanian = Moscovian-Gzhelian) in age (Stephenson et al., 2003). This indicates that it was laid down relatively early following formation of the Hercynian unconformity, and, further, that deposition commenced during an early stage of the late Paleozoic Gondwanan glaciation. It is clear that these fluvial sediments have a widespread distribution throughout the subsurface across the study area (Fig. 5). Similar extensive fluvial sands and gravels are identified within postglacial outwash deposits associated with the several phases of retreat of the Pleistocene ice sheets in northern Europe. It is inferred that the Unayzah C coarse-grained

45

sandstones and conglomerates analogously represent extensive glaciofluvial outwash deposits associated with retreat phases of the late Paleozoic Gondwanan ice sheets in Saudi Arabia. The occurrence of an extensive line of end-moraine complexes known as the Rehburg Line (Fig. 6) that is associated with the north European Pleistocene glacial outwash deposits is of special significance with regard to the Unayzah C member. This line extends >500 km from the North Sea in the west to Hanover, Germany, in the east (Bennett, 2001) and marks the approximate extent of glacier ice during the Rehburg Phase (Drenthe advance) of the Saalian Glaciation (Van der Wateren, 1995). Several push moraine complexes were documented along the Rehburg Line (Van der Wateren, 1985, 1987, 1994), wherein their architecture consists of a number of subhorizontal nappes that have been displaced horizontally by the ice, in some places as much as 6 km. These nappes are bounded above and below by a number of shear zones (Bennett, 2001). The description of the Unayzah C member presented above has highlighted the recognition of well-developed zones that show a range of deformation styles from brittle to ductile shear and between which the rock is essentially undeformed. Furthermore, recent investigations by the senior author of the Carboniferous-Permian lower Juwayl Formation (Unayzah C equivalent) at the outcrop in the Wajid region of southwest Saudi Arabia also revealed the presence of discrete zones of low-angle dislocation (including shear and overfolding) within sandstones that are interpreted as glacial outwash within glacial paleovalleys. (That outcrop work will be presented in detail in a future publication.) All of this evidence, considered in conjunction with the earlier conclusion that the top of the Unayzah C member represents an unconformity of some magnitude, suggests the following model for the evolution of the Unayzah C member. Following the mid-Carboniferous tectonic event and the near-coincident inception of the late Paleozoic Gondwanan South Polar glaciation, the ice ultimately extended across the southern part of the Arabian plate to the vicinity of wells 13 and 14 in this study. That is to say, it spread to paleolatitudes at least as far north as the present-day location of the Al Batin Arch (see Fig. 1A) and therefore significantly farther north than had been previously documented (e.g., Stampfli and Borel, 2004). This glacial event was long-lived and possibly persisted for some 30–45 m.y., from the mid-Carboniferous (late Visean to early Namurian, or Serpukhovian) to Early Permian (Sakmarian) times (Al-Husseini, 2004). Considerable uncertainty exists, however, regarding the specific timing of the onset of this glaciation in Arabia, not least because of the possibility of a number of separate phases of glacial advance and retreat within its overall duration. Melvin and Sprague (2006) and Osterloff et al. (2004a) discussed the likelihood that older glaciogenic deposits may very well have been successively removed as a result of extensive cannibalization during later glacial advances in Saudi Arabia and Oman, respectively. In the Unayzah C member in the subsurface of eastern-central Saudi Arabia, each of the inferred glacial retreat phases resulted in substantial volumes of glaciofluvial outwash sands and gravels

Figure 5. Stratigraphic section across the study area, showing thickness variations and lateral extent of the Unayzah C member, in relation to the Unayzah B member and the unnamed middle Unayzah member (“UmUm”). Datum is the top of the unnamed middle Unayzah member. Lower bounding surface is the Hercynian unconformity (HU). Note the extent of downcutting into the Unayzah B member by the pre-Khuff unconformity (PKU) in wells 2 and 3. Map shows line of section.

Figure 6. Map of the northwest European Plain, showing the Rehburg Line of push moraines. These moraines are associated with phases of glacial advance of the Pleistocene Fenno-Scandian ice sheet. Modified from Bennett (2001).

Late Paleozoic Gondwanan glaciation in Saudi Arabia being deposited across the area upon laterally extensive alluvial braided plains. During subsequent readvances the ice sheets progressed across the glaciofluvial outwash of each preceding retreat phase. In the process they formed a number of glacially induced, push moraine nappes, similar to those documented by Van der Wateren (1985, 1987, 1994) along the Rehburg Line from the Pleistocene of the north European Plain (Fig. 6). Ultimately, this process created a thick pile of subhorizontal, superimposed units of glaciofluvial sands and gravels, separated by distinct shear zones (Melvin and Sprague, 2006). It is likely that these repeated processes of glacial thrusting associated with readvances of the ice sheet were also responsible for the remobilization of pore fluids within the outwash sands, thus giving rise to the widespread evidence of dewatering that is observed within these rocks. The top Unayzah C unconformity is considered to represent the subglacial surface at the time of the final advance of the late Paleozoic Gondwanan ice sheet in Saudi Arabia. UNAYZAH B MEMBER The preceding discussion has shown how the Unayzah C member is characterized across the study area by a limited number of depositional facies, namely multistory quartzose sandstones and rare conglomerates that were deposited upon an extensive glaciofluvial outwash braided plain. Consequently, that member displays a fairly distinctive, monotonous wireline log signature. This is in stark contrast to the rocks that make up the Unayzah B member. Melvin and Sprague (2006) described in detail >1355 ft (406.5 m) of core from the Unayzah B member in 13 wells, and identified 8 distinctive depositional facies from within that stratigraphic unit. The characteristics of those various facies are presented below. In well 7, continuous core was recovered, not only from the entire Unayzah B member as it is represented in that well but also from the complete section through the overlying “un-named middle Unayzah member” (Melvin and Sprague, 2006), as well as the uppermost parts of the underlying Unayzah C member. This cored interval thus represents a crucial record of the stratigraphic relationships in the Unayzah in that well, and is reproduced as a reference section in Figure 7. Sediment Fracture-Fill Features (?Periglacial Deformation) In well 4, recovered cores show that the Unayzah B member rests directly upon the Hercynian unconformity (i.e., the Unayzah C member is absent at that location). A laminated sandstone 3 ft (0.9 m) thick is present ~2 ft (0.6 m) above the unconformity. It displays a sharp, subvertical fracture in core that is filled with a variety of coarse-grained detritus including mud clasts and granule-sized quartz grains. This feature has been discussed and tentatively interpreted by Melvin and Sprague (2006) as a frost contraction wedge formed in a periglacial setting. Similar features have been documented in modern periglacial settings (French,

47

1996; Ruegg, 1983), as well as from Proterozoic glaciogenic sediments from Greenland by Moncrieff and Hambrey (1990). Stratabound, Internally Deformed Deposits (Push Moraines) This facies is well displayed in two zones in well 7 (Fig. 7, 26.5–42.0 ft and 55–68 ft). At each location the rock displays severe structural deformation in the form of high- to low-angle reverse faulting (listric thrusting) (Fig. 8A) and overfolding. The rocks comprise sandstones and silty mudrocks that contain dispersed pebbles and small cobbles of siltstone and very fine– grained sandstone (diamictite) (Fig. 8B). Palynological analyses of these fine-grained deposits have yielded a heavily reworked assemblage of palynomorphs of Silurian, Devonian, and Early Carboniferous age, as well as significant quantities of monosaccate pollen that cannot be older than Namurian (Serpukhovian) in age (the “Cm palynoflora assemblage”: J. Filatoff, 2004, written commun.; Fig. 7). These intervals of deformed rock are stratabound by undeformed sediments laid down in a conformable succession (Fig. 7) and are thus interpreted as having been disrupted at, or very shortly after, their time of deposition. Furthermore, it can be shown (see below) that they are interstratified with sediments that are demonstrably glaciogenic in nature. They have thus been interpreted as the product of glacial deformation and specifically identified as the preserved remnants of glacial push moraines (Melvin and Sprague, 2006). Bennett (2001) described how push moraines display a wide range of morphologies at a range of scales, from a few meters to features that extend for several kilometers. The overthrust outwash sands and gravels described earlier from the Unayzah C member fall into the latter category. The stratabound, internally deformed deposits in the Unayzah B member are much smaller features, with an average thickness of ~14 ft (4.2 m). Boulton et al. (1999) identified four broad categories of push moraines, wherein the style of deformation involves either fans of imbricate thrusts or superimposed subhorizontal overthrusted nappes. Regarding the former, the smallest are generally no more than 16.4 ft (5 m) high and can be found in both terrestrial and subaqueous environments (Boulton, 1986; Boulton et al., 1999; Bennett, 2001). These small push moraines are more representative of the stratabound, internally deformed deposits seen in the Unayzah B member in well 7. In the western part of the study area the Unayzah B member in wells 2 and 3 (see Fig. 1A) directly overlies rocks of early Paleozoic age, i.e., it rests directly upon the Hercynian unconformity. It is 40 ft (12 m) thick in well 2, and 65 ft (19.5 m) thick in well 3, and in each well it is characterized throughout by poorly sorted sediment and by an abundance of small-scale faults and low-angle shear planes (Melvin and Sprague, 2006). This pervasive intraformational deformation of the Unayzah B member in these wells is also attributed to direct sustained contact with ice.

Figure 7. Core log from well 7, showing the complete section through the unnamed middle Unayzah member (“UmUm”) and the Unayzah B member in that well, and their respective contacts with the Unayzah A member above and the Unayzah C member beneath. Note: Cm indicates the locations of the proven occurrence of the heavily reworked “Cm palynofloral assemblage” in the Unayzah B member (J. Filatoff, 2004, written commun.).

Late Paleozoic Gondwanan glaciation in Saudi Arabia

49

Figure 8. Core photographs from well 7, showing features observed in the glacially tectonized facies of the Unayzah B member. (A) Interval of sandstone exhibiting high-angle, reverse faulted (thrusted) dislocations (see Fig. 7, 58–61 ft). (B) Poorly sorted pebbly sandstone (diamictite), displaying a strongly sheared fabric (see Fig. 7, 48–51 ft).

A

B

Massive, Very Poorly Sorted Pebbly Siltstones (Diamictite) Diamictites are poorly sorted sediments containing a wide range of particle sizes in a relatively fine matrix (Bates and Jackson, 1980). The massive diamictites described herein are commonly observed within the Unayzah B member in the Saudi Arabian subsurface, being recorded in this study from wells 1, 2, 3, 4, 5, 7, 10, and 12. In some cases they are severely deformed (e.g., wells 2, 3, and 7: see previous discussion) as a result of being incorporated in push moraines. Where they have been identified in core they range in thickness from ~6 ft (1.8 m) in well 10 to 14 ft (4.2 m) in well 12, to >200 ft (60 m) in well 5 (Fig. 9). In general the diamictites display almost no stratal fabric and are extremely poorly sorted. Thus they comprise a

host rock of gray-green (and locally red-brown) siltstone to very fine–grained sandstone within which occurs an abundance of well-rounded to subangular, dispersed grains of medium to very coarse quartz sand as well as up to 5% coarser material including granules and pebbles of granite, quartz, feldspar, black chert, gray siltstone, sandstone, and mudstone (Fig. 10A). Cobblesized clasts of fine-grained sandstone have been observed in places (Fig. 10B). Several workers investigating diamictitic deposits in glacial settings have identified a number of subfacies, including lodgment tillites, proximal and distal water-lain tillites, and debris flows (e.g., Visser, 1982; Levell et al., 1988; Moncrieff and Hambrey, 1990). The latter authors conceded that “there are numerous cases in which the interpretation remains open to question.” We

Figure 9. Core logs from wells 10, 12, and 5, showing associations of glaciolacustrine depositional facies in the Unayzah B member. (A) Sediment gravity flows in well 10 (0–25 ft; note the dropstone at 17 ft), overlain by multiple laminasets of ripple cross-laminated sandstones (25–44 ft) and a thin diamictite (44–50 ft), followed by mudrock in the uppermost 5 ft. (B) Laminated mudrock (0–7.5 ft) in well 12, overlain by massive diamictite (7.5–20.0 ft). This is followed by stratified diamictite (20–28 ft) and sediment gravity flows (28–41 ft) in the upper part. (C) Extremely thick interval of massive diamictite in well 5 on top of a unit of sediment gravity flows (0–10 ft). Wireline logs in this well suggest a total >200 ft (60 m) of diamictite.

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A

C

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D

concur with Melvin and Sprague (2006), who discussed in detail the massive diamictites of the Unayzah B member and concluded that they most likely represent resedimented glacial debris laid down at the bottom of glacial lakes as debris flow deposits. Interstratified Pebbly Siltstones and Laminated Mudstones: Stratified Diamictite This distinctive facies within the Unayzah B member has been recorded from wells 4, 7 (Fig. 7, 68–85 ft), and 12. It is heterolithic in character and comprises a host rock of very fine– grained, laminated mudstone (with associated rare, very thin beds of rippled, very fine–grained sandstone in some places), within which occur thin beds (millimeter to decimeter scale) of poorly

Figure 10. Core photographs showing features associated with massive and stratified diamictites in the Unayzah B member. (A) Typical example of massive diamictite, showing the very poorly sorted nature of the sediment, and the apparently random orientation of clasts (well 12). (B) Cobble-sized clast of fine-grained sandstone in massive diamictite (well 5). (C) Stratified diamictite in well 7, showing finegrained glaciolacustrine laminite with a thin (2 cm) interval of (type 1) very poorly sorted diamictite. Note the extreme fissility in the lower part of the laminite subfacies. (D) Cobble-sized clast of diamictitic material (till pellet) in type 1 stratified diamictite. Note the irregular, diffuse edges to the till pellet (well 12).

sorted pebbly siltstone material (diamictite) (Fig. 10C). Two types of this diamictite are observed. Type 1 beds are 0.2–4.0 in. (0.5–10.0 cm) thick and have indistinct upper and lower boundaries. They consist of silty mudstone that contains abundant dispersed grains of fine- to very coarse–grained sand, as well as rare granules of quartz, feldspar, and granite. Significantly, they also commonly contain distinctive pelletoid clasts of material that is itself very poorly sorted (diamictitic) in nature. These pelletoid clasts range in size from a few millimeters to 15 cm (Fig. 10D), are commonly flattened parallel to bedding, and show diffuse or “ragged” edges. Type 2 beds are 2–12 in. (5–30 cm) thick, with sharp, commonly loaded lower bed boundaries, and are similarly very poorly sorted. The most extensive development of this depositional facies recorded to date occurs in well 7 (Fig. 7, 68–85 ft).

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There, the type 1 beds become increasingly rarer toward the top of the 17 ft interval, where the laminated mudrocks predominate: this interval of stratified diamictite effectively fines upward in this well. The laminated mudrocks represent the background sedimentation of this depositional facies. To date, they have proved to be palynologically barren (N.P. Hooker, 2008, personal commun.) and are interpreted as lake-bottom deposits (as opposed to marine). In the type 1 diamictite beds the diffuse nature of the bed boundaries is suggestive of rain-out through the water column of debris derived from icebergs floating on the surface of the water body, probably during warm season meltout. The identification of the poorly sorted pelletoid clasts is significant in this regard, as they represent “till pellets” sensu Ovenshine (1970). That author considered such deposits to identify uniquely the existence of glacier ice in very close proximity to the environment of deposition. The type 2 diamictites, with their sharp bed contacts, more likely represent debris flow deposition on the bottom of glacial lakes. In well 7 the progressive diminution up-section of type 1 diamictite beds in favor of the laminated mudrocks (described above) suggests increasingly ice-distal conditions through time and/or progressive deepening of the lake. Nonstratified Pebbly Siltstone: Rain-Out Diamictite These rocks comprise a subfacies of the stratified diamictites, described above. In well 8, directly overlying the Unayzah C member, and constituting the entire Unayzah B member in this well, is a 2 ft (0.67 m) thick bed of diamictite. It is mud supported and very poorly sorted, and it displays no internal stratification. It contains an abundance of dispersed grains of fine- to very coarse–grained sand as well as rarer dispersed granules and small pebbles of quartz and sedimentary rock fragments. Melvin and Sprague (2006) illustrated how, at the top of this bed, there occurs a small (1 cm) pebble with its long axis oriented perpendicular to bedding. In the lower part of the bed, indistinct mottling is observed that is suggestive of bioturbation (Melvin and Sprague, 2006). Similar mottled diamictites are seen in well 7, within the stratified diamictite facies. The noted occurrence of the pebble with its long axis oriented perpendicular to bedding supports an origin for these rocks related to melt-out from floating ice in a glacial lake. Similar rain-out diamictites of glacial origin were described and discussed from the Karoo in South Africa by Visser (1982), albeit on a larger scale. The possible bioturbation in these Unayzah B sediments is suggestive of a seasonal aspect to their deposition and supported by its recognition in the stratified diamictites in well 7. Thus, in relatively warm periods, coarse-grained detritus was released from melting ice floes in a glaciolacustrine setting; having settled upon the lake bottom, the sediment was colonized by organisms that were suited to the relatively mild conditions. When harsher conditions prevailed, sediment rain-out was minimized, faunal activity dwindled, and sedimentation was reduced to suspension settling of very fine–grained silt and clay.

Multistory Graded Sandstones: Glaciolacustrine Gravity Flow Deposits This facies is recognized in core from wells 7, 9, 10, 12, 13, and 17. It is most fully developed in well 9, where it is >230 ft (69 m) thick. There it comprises sandstone beds that are 1–5 ft (0.3–1.5 m) thick, displaying sharp, commonly erosional basal contacts, and in many cases they are clearly graded (Fig. 11). They display an abundance of fluid escape structures, dominated by elutriation pillars and dish structures (cf. Lowe and LoPiccolo, 1974). These sandstones are organized into (at least) two upwardthinning and -fining bedset packages (Fig. 11). A similar package of sharp based, graded sandstones occurs in well 10 (Fig. 9A, 0–25 ft), and in that well a cobble sized, angular sandstone clast was recognized, “floating” on the top of one of the graded beds (Fig. 9A, 17 ft). Other occurrences of sharp-based graded sandstones in the Unayzah B member have been described in detail from other wells by Melvin and Sprague (2006). The graded sandstones described above from well 9 (Fig. 11) display all the characteristics of rapidly deposited, highly fluidized sediment gravity flows such as have been described by Lowe and LoPiccolo (1974). The organization of the beds into thick, upward-thinning and upward-fining packages is significant. It suggests sustained deposition in an environment within which, for each package of sediment, the deposits became increasingly distal with respect to their source. It is proposed that these gravity flow deposits represent sublacustrine outpouring of sediment from melting glacier ice. The two distinct upward-fining packages of sandstones may represent either the successive retreat of two different glacial sources (e.g., valley glaciers?), or they may reflect an element of glacial readvance and retreat in the location of this well. That these sublacustrine gravity flow sandstones in the Unayzah B member have a glacial origin is supported by the noted occurrence of the anomalously large clast described from well 10 (Fig. 9A). This is interpreted as a dropstone, which melted out of ice floating on the lake surface, and fell upon the lakebottom gravity flow beneath. In well 13, dropstones have been recognized in fine-grained lake deposits (Melvin and Sprague, 2006) that were subsequently overlain by upward-thinning and -fining packages of gravity flow deposits, similar to the deposits described herein at well 9. Furthermore, the overall trend of those deposits in well 13 was to become thinner and finer upward, suggesting that the sediment supply was becoming more distant (i.e., the rocks became more ice-distal in character), and/or the lake was becoming deeper. Ripple-Drift Cross-Laminated Sandstone: Sublacustrine Glacial Outwash This depositional facies has been observed only in core from well 10, where it occurs in a unit that is ~20 ft (6 m) thick (Fig. 9A, 25–45 ft), resting directly above the interval of sublacustrine gravity flow sandstones mentioned above. It comprises three bedsets, 3–11 ft (0.9–3.3 m) thick, each separated from the other

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Figure 11. Wireline gamma-ray (GR) log, showing cored interval and associated core log from well 9, showing thick development of multiple beds of highly fluidized sediment gravity flows (as indicated by abundance of dish structure and elutriation pipes). Note how the beds are arranged in thick packages that show a general upward thinning and fining trend (see text for full discussion). These gravity flow sandstones sit directly upon an interval (uncored) with a high gammaray-log signature, which has yielded samples of the heavily reworked Cm palynoflora (J. Filatoff, 2004, written commun.).

by a thin interval of siltstone. Each bedset consists of fine- to very fine–grained sandstone characterized by multiple, thin (1–3 cm) laminasets of climbing ripple-drift cross-lamination. Type A ripple-drift cross-lamination of Jopling and Walker (1968) dominates the succession, although it is seen to pass upward into type B ripple-drift cross-lamination in the uppermost 3 ft (0.9 m). The sandstones that display type A ripple-drift crosslamination have been interpreted to represent sustained bedload transport of sediment that originated as subaqueous outwash material in a glacial lake (Melvin and Sprague, 2006). Analogous sublacustrine glacial outwash deposits, dominated by laminasets of ripple-drift cross-laminated sandstones, have been described from the Pleistocene of Canada, where they are associated with lake-bottom kame deltas (Gustavson et al., 1975) as well as subaqueous esker fan deposits (Rust and Romanelli, 1975). The uppermost sets of type B ripple-drift cross-lamination in well 10 in the present study suggest decreasing energy, whereby increased fallout from suspended load prevailed over bedload transport. This inferred loss of energy is taken to indicate that the original source of sediment (melting ice) had become somewhat farther removed from the location of well 10, i.e., this succession of ripple-drift cross-laminated sandstones represents an increasingly ice-distal facies association. Mudrock Facies: Distal Glaciolacustrine Deposits Fine-grained sediments of this depositional facies have been observed in a number of wells (e.g., wells 7, 10, 12, 13, and 17) (Melvin and Sprague, 2006). In well 10 (Fig. 9A, 50.5–69.5 ft), 19 ft (5.7 m) of dark gray mudrock overlies a 6 ft (1.8 m) thick diamictite. The lowermost 10 ft (3 m) of this mudrock deposit comprises a series of stacked, thin (centimeter scale) muddy siltstone beds that are sharp based and are distinguished from each other only by very thin (millimeter scale) claystone partings. This lower unit fines upward gradationally into increasingly argillaceous, homogeneous dark gray mudrock; in the uppermost 0.13 ft (4 cm) it becomes a pale gray claystone. The muddy siltstones of the lower part of this mud-prone interval display the characteristics of very fine–grained gravity flow deposits (“distal turbidites”) and pass upward into essentially featureless mudstone, interpreted as suspension fallout deposits. They have been proven to be palynologically barren (J. Filatoff, 2004, written commun.) and are considered to have been laid

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down on the bottom of a lake. The overall upward-fining of these muddy sediments from siltstone to claystone is considered to represent a gradual deepening of the lake through time. Origin and Evolution of the Unayzah B Member The foregoing discussion has highlighted the large number of depositional facies that characterize the Unayzah B member. Except for the inferred periglacial frost wedging described earlier, and the case of wells 2 and 3 in the western part of the study area (to be discussed below), each of those facies is interpreted to have been deposited in a lacustrine environment (palynological evidence for a marine setting being consistently lacking). In most cases there is evidence, direct or indirect, to suggest a glacial influence upon that environment. Thus, the stratabound, internally deformed deposits in well 7 have been interpreted above to represent the remains of glacial push moraines and as such provide strong evidence of an ice-contact setting. The remaining depositional facies are indicative of a spectrum of glacially related environments, ranging from ice-proximal to ice-distal. Considering the volumes of sediment they represent, the >200-ftthick, massive, very poorly sorted pebbly siltstones (diamictites) at well 5 (Fig. 9C), and the >230-ft-thick sequence of multistory graded sandstones (glaciolacustrine gravity flow deposits) in well 9 (Fig. 11) most likely represent deposition in an ice-proximal sublacustrine environment. In the latter case the organization of the gravity flow deposits may suggest either multiple glacial sources or possibly some degree of glacial readvance, as has been discussed earlier. Ice-distal settings are suggested by the very fine–grained sediment (mudrock facies) seen in well 10 as well as by the interstratified pebbly siltstones and laminated mudstones (stratified diamictite) that are developed in wells 7 and 12. Sediments indicative of deposition in settings intermediate in the “proximal-distal spectrum” are represented by nonstratified pebbly siltstone (rain-out diamictite) and ripple-drift crosslaminated sandstone (sublacustrine glacial outwash), as well as the thinner bedded occurrences of massive, very poorly sorted pebbly siltstones (diamictite) and multistory graded sandstones (glaciolacustrine gravity flow deposits). The assignment of these various depositional facies to a position in the “ice proximal-distal spectrum” is substantiated, not only by their grain size, thickness, and general sedimentological characteristics, but also by their associations relative to each other from well to well. That is to say, the vertical association of depositional facies commonly suggests an overall passage in any given location from an ice-proximal setting to one that was ice-distal. This is well displayed in well 10 (Fig. 9A). There, a lowermost interval of at least 25 ft (7.5 m) of medium- to coarsegrained multistory graded sandstones (glaciolacustrine gravity flow deposits) (with dropstones) represents ice-proximal sublacustrine fan deposition. This is overlain by ~20 ft (6 m) of multiple laminasets of ripple-drift cross-laminated sandstone (Fig. 9A, 25–45 ft). These sediments were interpreted earlier also to be sublacustrine glacial outwash detritus, within which the upward pas-

sage from type A to type B ripple-drift cross-lamination (sensu Jopling and Walker, 1968) suggests increasingly ice-distal conditions. They are directly overlain by a massive diamictite deposit that is only ~6 ft (1.8 m) thick. Its limited thickness suggests an intermediate setting in terms of its proximity to any glacial source (especially when compared with the very great thickness of the same facies at well 5). The uppermost 19 ft (5.7 m) of the Unayzah B succession in well 10 comprises ~10 ft (3 m) of mudprone distal turbidites that pass up into 9 ft (2.7 m) of featureless mudrock (Fig. 9A, 50.5–69.5 ft). These fine-grained sediments were laid down in an extremely ice-distal setting. This sedimentary succession at well 10 thus represents a passage through time from a relatively ice-proximal setting to one that was distinctly ice-distal. The clear inference from this evidence is that, relative to the location of well 10, either the ice was becoming increasingly distant (i.e., retreating) or the lake was becoming deeper, or both. It was noted above how Melvin and Sprague (2006) observed similarly that in well 13, a series of upward-thinning and fining packages of gravity flow deposits in the Unayzah B member displayed an overall trend toward being thinner and finer upward, again suggesting that the sediment supply (from the ice) was becoming more distant or the lake was becoming deeper. In well 7 the sediments of the Unayzah B member that directly overlie the top Unayzah C unconformity consist of ~23 ft (6.9 m) of sublacustrine gravity flow deposits (Fig. 7, 3–26.5 ft). These are interpreted as relatively ice-proximal, sublacustrine fan deposits. They are overlain (Fig. 7, 26.5–42.0 ft) by the lower of two units of stratabound, internally deformed deposits (push moraines), inferred to represent ice-contact conditions. The succession continues upward with ~13 ft (3.9 m) of massive, very poorly sorted pebbly siltstones (diamictite) (Fig. 7, 42.0–55.0 ft), and that is again interpreted to be a relatively ice-proximal deposit. The return of an ice-contact setting is seen in the second of the stratabound, internally deformed deposits (push moraines) (Fig. 7, 55–68 ft). The highest part of the Unayzah B member in well 7 (Fig. 7, 68–85 ft) is characterized by well-developed and sustained deposition of interstratified pebbly siltstones and laminated mudstones (stratified diamictite). Those mud-prone sediments represent a significantly more ice-distal setting than any of the underlying sediments. Furthermore, the upward diminution of the amounts of glacially derived sand grains in this facies in this well (described earlier) does itself strongly imply continued retreat of the ice from this location. Thus, in well 7 the Unayzah B member displays evidence for an overall gradual retreat of the ice whereby ice-proximal conditions gave way to ice-distal conditions, but during which at least two distinct icecontact deformational events occurred. That is to say, the overall retreat of the ice was interrupted by at least two minor readvances of the ice sheet in this location. The possibility of minor glacial readvance within the Unayzah B member was also inferred earlier from the architecture of the gravity flow sandstones described from well 9. It is clear from the foregoing examples from individual wells that across most of the study area the vertical association of

Late Paleozoic Gondwanan glaciation in Saudi Arabia depositional facies in the Unayzah B member generally displays an increasingly ice-distal aspect up-section. In all cases this is interpreted as representing the steady retreat of the Gondwanan ice sheet in early Permian times. Locally, there is evidence for some minor readvance of the ice (e.g., at wells 7 and 9). Support for these sedimentological conclusions can be found in the palynology. The Late Carboniferous–Early Permian assemblages that characterize the OSPZ2 palynozone of Stephenson et al. (2003) (which in turn characterizes the Unayzah B member in Saudi Arabia) comprise a variety of spores, pollen, and fresh-water algae (N.P. Hooker, 2008, written commun.). These assemblages can be interpreted as showing a progression from a cold, dry climate in the Late Carboniferous to a slightly warmer and wetter climate in the Early Permian. The latter represents the transition from a glacial to a deglacial phase, and is reflected in the palynoflora by the change from low diversity spores and monosaccate pollen-dominated assemblages (glacial phase) to high diversity spores with fresh-water algae (deglacial phase) (N.P. Hooker, 2008, written commun.). The palynofloral assemblages of the deglacial phase are commonly derived from the massive diamictites described above as being relatively common within the Unayzah B member. Figure 12 is a stratigraphic section that is hung using the top of the “un-named middle Unayzah member” (Melvin and Sprague, 2006) as a datum. It incorporates both the Unayzah B member and the “un-named middle Unayzah member.” It is clear that the Unayzah B member exhibits considerable thickness variation across the study area. Specifically the thickest sections of this stratigraphic unit are seen in wells 5 and 9: it has been discussed above how these two wells contain sediments that are the most ice-proximal in their character. The thinnest occurrence of the Unayzah B member is at well 8 (Fig. 12), where it is represented by only 2 ft (0.6 m) of nonstratified pebbly siltstone (rain-out diamictite) resting directly upon the top of the Unayzah C unconformity. Considered in the context of the preceding discussion, these stratigraphic relationships suggest the following model for the evolution of the Unayzah B member. Following the various retreats and inferred readvances of the Gondwanan ice sheets that led to the deposition and deformation, respectively, of the Unayzah C member (see earlier discussion), the subglacial surface of the final readvance was effectively preserved as the top of the Unayzah C unconformity. Thereafter, the final melting and ultimate withdrawal of the ice took place. As the ice melted, the subglacial topography was exposed, and the lows began to be filled with the meltwaters from the retreating ice sheets, with the consequent formation of abundant glacial lakes. Initially the deepest of those lakes were infilled with large volumes of resedimented glacial outwash material in the form of massive, very poorly sorted ice-proximal debris flows (diamictites) (e.g., well 5, Fig. 9C) and highly fluidized gravity flow sands that developed significant lake bottom turbidite deposits (e.g., well 9, Fig. 12). As the ice continued to retreat, more meltwater and more sediment were supplied, and the numerous glacial hollows were filled and spilled over with rising floodwaters.

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Locally minor readvances of the ice occurred, with the creation of minor push moraine deposits. In general, however, the terminal retreat of the ice was reflected and preserved in the sedimentary record in the form of increasingly ice-distal deposits. The landscape changed from one of numerous isolated and deep glacial lakes to one that was awash with glacial meltwaters as the lakes overspilled their boundaries and became connected. This terminal (maximum) flood scenario is preserved at well 8, where the thin rain-out diamictite directly overlies the top of the Unayzah C unconformity and appears to link the much deeper glacial lake deposits that characterize wells 7 and 9 (Fig. 12). In the western part of the study area, around wells 2 and 3 (Fig. 12) a somewhat different situation prevailed. It was noted earlier how in each of these two wells the Unayzah B member is represented in its entirety by relatively thick intervals of glacially deformed sediment, which are attributed to sustained, direct contact with the ice. Furthermore, it is clear from the lack of relevant depositional facies that when the ice did eventually melt in this area, the meltwaters were not ponded in topographic lows to form significant lakes such as has been described widely from across the region. The possible corollary exists therefore that the western part of the study area may have been topographically high during Unayzah B times. This is a not altogether unlikely scenario, given the proximity of these westerly locations to the Al Batin Arch of Faqira et al. (2009) (see Fig. 1A). Indeed, given the suggestion by some workers (e.g., Eyles, 1993) that ice coverage may have extended over to present-day east Africa, it is interesting to speculate upon the possibility of a center of high altitude (alpine) glaciation in western central Saudi Arabia at this time. UNNAMED MIDDLE UNAYZAH MEMBER The unnamed middle Unayzah member was recognized and defined in the extensive core-based study of the lower Unayzah Formation carried out by Melvin and Sprague (2006). Its facies characteristics show that it is very different from the glaciogenic Unayzah B member, and yet it displays a distinctive character that sets it aside also from the Unayzah A member. It is a stratal unit that represents a transition from the Unayzah B member to the Unayzah A member, even though its boundaries are sharp and readily identified in core. Among the wells examined in this study, the unnamed middle Unayzah member occurs in a number of situations. Thus it may be identified directly overlying the Unayzah C member, as at well 6 (Fig. 12), although more commonly it sits upon the glaciogenic facies of the Unayzah B member. In places it is absent, as at well 11 (Fig. 12), and elsewhere it may have been removed by erosion at the pre-Khuff unconformity, as seen in wells 2 and 3 (Fig. 12). In well 7 the unnamed middle Unayzah member appears to have been completely cored, and its relationships to the Unayzah B and A members are clearly exposed (Fig. 7). This well thus serves as a valuable reference well. The nature of the individual depositional facies and the stratigraphic boundaries of the unnamed middle Unayzah member are discussed below.

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Figure 12. Stratigraphic section across the study area, showing facies architecture within the Unayzah B member and the overlying unnamed middle Unayzah member. Datum is the top of the unnamed middle Unayzah member. In wells where this sediment package is thick (e.g., wells 7, 9), the lowermost parts of the section are dominated by glaciogenic deposits (Unayzah B member). These deposits are overlain in those wells by nonglacial, red lacustrine siltstones and associated sandstones of the floodplain facies of the unnamed middle Unayzah member. Where the overall package is thin (e.g., well 6), it comprises only the latter member. Note the effects of the pre-Khuff unconformity (PKU) in wells 2 and 3. Map shows line of section. Modified from Melvin and Sprague (2006). GR—gamma-ray.

Fluvial Sandstones The best example of fluvial sandstones in the unnamed middle Unayzah member is seen in core from well 8. There, Melvin and Sprague (2006) identified a sandstone interval that is 30.5 ft (9.15 m) thick (Fig. 13A, 29.5–60.0 ft). It directly overlies some 10 ft (3 m) of red sandy siltstones (alluvial floodplain deposits: see below) and occurs ~18 ft (5.4 m) above the top of the very thin Unayzah B member in this well. It is also overlain by a thick interval of 45 ft (13.5 m) of red sandy siltstones of the alluvial floodplain facies (see below). Fine- to mediumgrained, moderately sorted sandstones occur in stacked beds that are 2–4 ft (0.6–1.2 m) thick. They display sharp, erosional contacts, with local mud-clast concentrations, and pass upward into low-angle trough cross-lamination overlain in places by finer–

grained, ripple cross-laminated sandstones (Fig. 13A). These characteristics of the sandstones, and their occurrence embedded within red very fine–grained (silty) deposits, suggest deposition in a channelized, heavily oxidized (i.e., terrestrial) setting. They are thus interpreted to be fluvial in origin, and their architecture of stacked shallow channel deposits encased and isolated within finer–grained floodplain deposits further suggests deposition in relatively high-sinuosity or even anastomosing rivers. Fluvial channel sandstone deposits, isolated within fine-grained floodplain siltstones, were described from the Cutler Formation (Permian to Pennsylvanian) of New Mexico by Eberth and Miall (1991), and were similarly interpreted to be anastomosed river deposits. In well 7 there occurs a similar interval of sandstones that shows cross-bedding and an upward-fining profile; these are also considered to be of fluvial origin (Fig. 13B, 41–50 ft).

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Figure 13. Core logs from wells 8, 7, and 6, illustrating the nature of the unnamed middle Unayzah member (“UmUm”). Note in particular the distinct sandstone units overlying siltstone intervals, and the similarities and differences displayed among them. See text for detailed discussion.

Eolian Sandstones This facies is well displayed in the unnamed middle Unayzah member in wells 7 and 6 (Fig. 13B and 13C). In well 7, there is an interval of ~17 ft (5.1 m) of low-angle to flat-laminated, well-sorted fine-grained sandstones wherein the grains are well rounded and appear frosted (Fig. 13B, 53–70 ft). In well 6, texturally similar flat-laminated sandstones are interbedded with more poorly sorted sandstones, displaying well-developed adhesion ripples (Fig. 13C, 23–40 ft). Thin (centimeter scale) beds of

dark gray siltstone are also present. Overlying these sandstones in well 6 is an interval 5 ft (1.5 m) thick comprising well-sorted, well-rounded, and frosted grains that occur in high-angle, grainsize-segregated cross-laminations (Fig. 13C, 40–45 ft). The cross-laminations are disrupted by small-scale synsedimentary faults (Fig. 13C, 42–43 ft). Synsedimentary deformation in similar facies in the unnamed middle Unayzah member has been observed in several other places. It is particularly well developed in western parts of the study area in central Saudi Arabia, in the vicinity of well 24 (see Fig. 1A).

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Melvin and Sprague (2006) described how the base of the unnamed middle Unayzah member in well 7 features a thin (2 ft; 0.6 m) red sandstone that rests abruptly upon the dark gray stratified diamictite at the top of the Unayzah B member in that well (Fig. 13B, 4–6 ft). That sandstone has high-angle crosslaminations that become quite diffuse in the lower levels of the bed. It also displays local nested concentrations of subrounded to subangular, granule-sized grains. It has a sharp, granule- and pebble-strewn basal contact with the underlying glaciolacustrine stratified diamictites of the Unayzah B member. The latter rocks display considerable cracking in their uppermost part, and the red sandstone is seen to penetrate deeply into those dislocations (Fig. 14A). The flat-laminated sandstones described above from well 6 and well 7 have the character of eolian sand sheet deposits, such as were well documented by Fryberger et al. (1979). The

A

beds showing adhesion ripples represent damper, sandy interdune conditions, and the thin siltstone interbeds suggest shortlived, shallow bodies of water in an otherwise eolian dominated environment. The overlying cross-laminated sandstone interval in well 6 is interpreted to be a residual eolian dune deposit. The synsedimentary dislocations may represent gravitational collapse on the dune slip face. Alternatively, Melvin and Sprague (2006) suggested the possibility that these dislocations may relate to sediment readjustment in response to melting of bodies of ice or snow that may have been trapped during dune migration. Analogous features were described from modern cold-climate eolianites from the Rocky Mountains of Colorado by Ahlbrandt and Andrews (1978) and from coeval Permian deposits of the Gondwanan sequence of Australia (Williams et al., 1985). Ostensibly, the thin red sandstone that occurs at the base of the unnamed middle Unayzah member in well 7 (Fig. 13B,

B

Figure 14. Core photographs showing the nature of the lower and upper contacts of the unnamed middle Unayzah member in well 7. (A) Gray, silty glaciolacustrine mudstone at the top of the Unayzah B member, abruptly overlain by red eolian sandstones at the base of the unnamed middle Unayzah member (“UmUm”) (see Fig. 13B, 5 ft). At the contact, note how the glaciolacustrine deposits are cracked, and the coarse basal sediment of the red sandstones has deeply invaded the cracks. (B) Red siltstone and silty, very fine–grained sandstone. The mottled and rooted lower interval is the uppermost part of the unnamed middle Unayzah member (“UmUm”) in this well, and represents a paleosol (Fig. 13B, 72.5–77 ft). The abruptly supradjacent siltstones are assigned to the Unayzah A member.

Late Paleozoic Gondwanan glaciation in Saudi Arabia 4–6 ft) displays the character of a residual, basal eolian dune sandstone deposit. The juxtaposition of such terrestrial sediment relative to the underlying glaciolacustrine stratified diamictites has significant stratigraphic implications that are discussed further below. Red Sandy Siltstones: Alluvial Floodplain Deposits The unnamed middle Unayzah member is commonly characterized by extensive deposits of sandy siltstones and silty, very fine–grained sandstones that are generally red, ranging from reddish brown to reddish purple, and locally red-green variegated. They are well displayed in core from well 6 (Fig. 13C, 10.5– 22.0 ft) and well 7 (Fig. 13B, 7–33 ft). In well 8, 10 ft (3 m) of red siltstones are seen below the fluvial sandstones described above (Fig. 15, 19.5–29.5 ft), and a further 47 ft (14.1 m) of red sandy siltstones are present above those sandstones (Fig. 15, 60–107 ft). In that well the highest parts of the unnamed middle Unayzah member show a distinct, albeit subtle upward-coarsening grain size trend (Fig. 15, 107–129 ft). Although generally these red sandy siltstones display very diffuse bedding characteristics, in places they are interbedded with red, silty very fine–grained sandstones that can display irregular lamination or occur as very thin (centimeter scale), current-rippled beds. Rarely, these beds have sharp upper contacts, showing evidence of sand-filled desiccation cracks. These fine-grained red beds were laid down under highly oxidizing conditions, as suggested by the widespread red coloration. They are interpreted to have been deposited in a terrestrial environment dominated by low energy conditions. The generally diffuse bedding characteristics are indicative of heavily waterlogged sediment, and the local development of ripple lamination suggests the presence of minor current activity. Thus this facies is considered to represent subaqueous deposition within shallow floodplain lakes. Locally, these lakes were at times infilled to the point of exposure and desiccation. In that context, the upwardcoarsening interval at the top of the unnamed middle Unayzah member in well 8 (Fig. 15, 107–129 ft) is tentatively interpreted as a small floodplain lacustrine delta. Paleosols In well 7, the uppermost part of the unnamed middle Unayzah member is represented by ~4 ft (1.2 m) of red-purple-gray variegated argillaceous and silty very fine–grained sandstones (Fig. 13B, 72–76 ft; Fig. 14B). These sandstones are tightly silica cemented and poorly sorted, containing an abundance of dispersed (floating) grains of medium- to coarse-grained sand. They are characterized by well-developed root traces up to 1 ft (30 cm) long, cutans, and other indicators of pedogenic development (Fig. 14B). In well 6 the eolian sandstone described earlier from the unnamed middle Unayzah member is overlain by ~4 ft (1.2 m) of poorly sorted sediment (Fig. 13C, 46–50 ft) that contains abundant floating grains of medium to coarse sand. In many

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places these sand grains appear to be concentrated within subvertical cracks within the sediment. This interval in well 6 displays characteristic subspherical fractures throughout. Similar features have also been noted at the top of the unnamed middle Unayzah member in well 9. These variegated, texturally immature rocks are interpreted as paleosols. In most places where they are recognized, they represent the stratigraphically highest occurring facies within the unnamed middle Unayzah member. Origin and Evolution of the Unnamed Middle Unayzah Member The contact that separates the Unayzah B member from the unnamed middle Unayzah member is extremely sharp (Fig. 14A) and is recognized across the study area (Fig. 12). Melvin and Sprague (2006) interpreted the contact to represent a regionally disconformable surface. The various facies that characterize the uppermost part of the underlying Unayzah B member represent widespread lake-dominated sedimentation at the time of maximum flooding of glacial meltwaters across the landscape in Early Permian times, as discussed above. The (possible) cold-climate eolian, fluvial, and fine-grained alluvial red-bed deposits of the unnamed middle Unayzah member clearly represent a very different, terrestrially dominated depositional setting and suggest that a highly significant drainage event marked the end of Unayzah B deposition. Martini and Brookfield (1995) described how, in the Pleistocene Bowmanville Bluffs in Ontario, Canada, a very sharp demarcation occurs between clay-rich glaciolacustrine rhythmites and overlying sand-prone sediments. They suggested that the widespread distribution of this contact and the nature of the overlying successions were the result not of erosion but of a sudden drop in lake level caused by the retreat of a glacier. Teller et al. (2002) noted that during the last deglaciation of the Quaternary period, melting of the Laurentide Ice Sheet in North America led to the release of very large volumes of stored precipitation to the oceans. There were a number of lakes along the margin of the Laurentide Ice Sheet, whose confining ice or sediment dams failed during the deglaciation. Of these, proglacial Lake Agassiz was by far the largest, covering a total of more than one million square kilometers over its 4000 yr history (Teller et al., 2002). Frequent changes in lake levels of Lake Agassiz have been documented. These changes were abrupt and often involved the release of several thousand cubic kilometers of water. The exact time for each lake drawdown (outburst) can only be estimated, but it is believed that lake levels for most phases could have been drawn down in a matter of months to a few years (Teller et al., 2002). Similarly, in northern Russia during the last glaciation of the Quaternary Period, the North Polar ice sheets expanded and blocked north-flowing rivers such as the Yenissei, Ob, and Pechora. As a result, south of the ice sheets a number of large ice-dammed lakes formed that were considerably larger than any lake on Earth today (Mangerud et al., 2004). The final

Figure 15. Core log through ~240 ft (72 m) of core in well 8, extending from the uppermost part of the Unayzah C member to the top of the Unayzah A member. Note the stratigraphic boundaries and the overall upward-coarsening profile from 132 to 238 ft. Modified from Melvin et al. (2005).

Late Paleozoic Gondwanan glaciation in Saudi Arabia drainage of the best mapped lake (namely Lake Komi in the Pechora Lowlands of northern Siberia) was modeled, and it was concluded that it probably emptied within a few months (Mangerud et al., 2004). These results dramatically demonstrate the likelihood of geologically instantaneous drainage events related to glaciolacustrine settings at times of terminal glacial retreat. Following the probably dramatic drainage event that is inferred to have terminated glaciolacustrine deposition of the Unayzah B member, the Permian landscape was dominated for the most part by very low lying alluvial floodplains that accommodated deposition of the large volumes of fine-grained (silt-sized) material that now constitutes the greater part of the unnamed middle Unayzah member. A number of river systems traversed these flood basins, but they were probably relatively isolated in occurrence, as is indicated from the limited evidence in the apparent relative isolation of their resulting channel deposits (Figs. 13A and 15). Toward the end of the period of deposition of the unnamed middle Unayzah member, relatively drier conditions became established, accompanied by a reduction in sediment supply, and the fluvial sands were reworked into a number of isolated eolian deposits (Figs. 13B and 13C). The possibility that these sandstones may represent cold-climate eolianites has considerable implications for the Permian paleogeography of the study area, particularly in the western part. There, around well 24 (Fig. 1A), these postulated cold-climate eolian deposits are particularly well developed. It was mooted earlier that the possibility exists that a high-altitude Permian glaciation was active and centered on the Al Batin Arch (see Fig. 1A). If such was the case the likelihood that these deformed eolian sandstones of the unnamed middle Unayzah member are indeed cold-climate deposits, laid down at a relatively high altitude, is rendered significantly more credible. Significantly, ongoing palynological studies from the same general area appear also to suggest assemblages having “montane” affinities (N.P. Hooker, 2008, personal commun.). Across the study area as a whole, active deposition of the unnamed middle Unayzah member ultimately ceased, and soilforming processes prevailed. These resulted in the formation of the paleosols that in many places characterize the top of the unnamed middle Unayzah member. Those soils represent a depositional hiatus and consequently mark a disconformable surface that separates the unnamed middle Unayzah member from the overlying Unayzah A member. UNAYZAH A MEMBER The Unayzah A member has a widespread distribution across the study area. Generally its thickness ranges from 150 to 300 ft (45–90 m) but does not display the extreme variation recognized in the Unayzah C or Unayzah B members. Although identified in the subsurface across the area of investigation, the Unayzah A member nonetheless is difficult to correlate in detail from wireline logs. This is because of the intrinsic variability in depositional facies distribution, confirmed from examination of >3150 ft (945 m) of Unayzah A core in the present study. This

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work has identified several different depositional facies (summarized below), and these and their resultant facies associations have permitted the recognition of three major depositional environments within the Unayzah A member. All of these environments are strongly indicative of a highly arid continental setting, and hence are suggestive of a significant and sustained climate change that became manifest at the beginning of deposition of the Unayzah A member. Semiarid Ephemeral Lake Deposits These rocks characterize the lower intervals of the Unayzah A member almost everywhere it occurs. They comprise silty, very fine–grained sandstones with minor sandy siltstones that range in color from brick red and red brown to buff yellow and pale to dark gray. They display variably lenticular to indistinct crinkly lamination, and locally small-scale ripple cross-lamination is seen. Rarely, over-steepened and folded sediment is observed. Elsewhere, small-scale vertical features have been noted and interpreted as syneresis cracks. These very fine–grained rocks are specifically characterized by the common occurrence of thin, flat-lying (horizontal) laminae, commonly only one or two grains thick (Fig. 15, 136–186 ft). The laminae comprise grains of wellrounded, medium to coarse sand. Commonly, bedding and lamination within these rocks are disrupted by subvertical, sand-filled cracks that have a downward penetration of 2–5 cm. These fine-grained silty sediments dominate the lower parts of the Unayzah A member, as was noted above. As such, they commonly overlie the similar red siltstones that were described earlier from the unnamed middle Unayzah member, separated from them in most places only by the paleosols that mark the end of deposition of the latter unit. The subtle but critical difference between these two siltstone facies lies in the abundance of coarser-grained sand laminae, as well as a greater number of sand-filled vertical cracks, that are observed within the lower zones of the Unayzah A member. The siltstones within the unnamed middle Unayzah member have been interpreted above as the depositional infill of alluvial floodplain lakes. Sedimentological evidence for those lakes being relatively long-lived bodies of standing water is much greater than is seen in the lower Unayzah A member. In the latter case, the sediment is similarly interpreted to be of a shallow lacustrine origin, wherein waterlain deposits are indeed preserved (as indicated by the ripple forms and evidence of slumping). However, there is much more abundant evidence for fluctuation of the lake levels, and consequent repeated and widespread exposure and desiccation within these lower Unayzah A silty, very fine– grained sandstones. That evidence includes the common occurrence of subvertical sand-filled cracks (desiccation cracks) and the abundant horizontal laminae of coarser-grained sand. These features are interpreted as being related to quasi-planar adhesion laminae (sensu Hunter, 1980) or adhesion laminations (Kocurek and Fielder, 1982) that were laid down upon exposed damp surfaces. The lowermost interval of the Unayzah A member is thus

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interpreted to represent deposition within an ephemeral lake basin setting (playa lakes) and as such indicates much drier conditions than those that prevailed during deposition of the unnamed middle Unayzah member. Semiarid Ephemeral Stream Deposits These deposits constitute the greater part of the upper Unayzah A member in a number of wells across the study area. Thus, for example in well 8, the upper part of the Unayzah A member displays a number of interbedded sandstones and siltstones (Fig. 15, 186–238 ft). The sandstones are thin-bedded units, rarely more than ~1 ft (30 cm) thick, that make up the lower parts of upwardfining couplets. They are fine to medium grained, moderately to poorly sorted, and generally either massive or flat laminated with sharp, erosional lower contacts (locally with numerous small mud clasts). These sandstones can be single units, or in places they comprise thin amalgamated beds (Fig. 16A). They are commonly overlain by finer-grained beds of similar thickness that form the upper parts of the depositional couplets referred to above. These beds comprise very fine–grained silty sandstone that displays well-developed ripple cross-lamination (Fig. 16B). In places these silty, rippled sandstones contain an abundance of mud drapes that display well-developed desiccation features (Fig. 15, 194–196 ft; Fig. 16C). The lower, sharp-based and coarser sandstones of these upward-fining couplets are interpreted as having been deposited rapidly by flash floods, and the overlying, finer-grained, and ripplelaminated sediment represents the waning flow deposits of those

A

B

flood events. The repeated occurrence of desiccation in the finer, mud-draped sediment emphasizes the episodic nature of deposition of these beds in a semiarid environment; i.e., these are the sporadic deposits of ephemeral streams. The overall depositional profile of the Unayzah A member at well 8 (Fig. 15, 132–238 ft) displays a clear upward-coarsening trend, passing gradationally from a siltstone-dominated regime upward into medium- and coarse-grained sandstones. The siltstone-dominated interval represents the semiarid ephemeral lake (playa lake) deposits of the lower Unayzah A member, discussed above. The upward-coarsening depositional trend that overlies these playa lake sediments in well 8 represents the progradation into the lake of a terminal alluvial fan, or a lake-marginal bajada, that was traversed by a number of shallow, ephemeral stream channels. Eolian Erg Center to Erg Margin Deposits Although the semiarid ephemeral stream deposits described above are encountered in the upper Unayzah A member in several places in the subsurface of eastern and central Saudi Arabia, they do not everywhere represent the dominant depositional facies association of that stratigraphic unit. In many places the uppermost part of the Unayzah A member consists of a complex of sandstone facies that can be ascribed in general to an eolian setting (Fig. 17), and which are summarized below. Eolian dune cross-bedded sandstones are common throughout the upper Unayzah A member. They comprise fineto medium-grained, well to very well sorted sandstones, with very well rounded and frosted grains of quartz sand that occur

C

Figure 16. Core photographs illustrating aspects of the semiarid ephemeral stream facies association of the upper Unayzah A member (see text for discussion). (A) Amalgamated medium-grained sandstones representing flash flood deposition. Note the sharp, scoured, and loaded contacts (arrows), suggestive of rapid, high energy sedimentation. (B) Argillaceous, ripple laminated fine- to very fine–grained sandstones indicative of waning flow conditions following flash flood events. (C) Abundance of desiccated clay drapes in the ripple laminated facies (arrows).

Figure 17. Core log through ~210 ft (63 m) of core in well 21, representing the various facies identified in the eolian facies association of the upper Unayzah A member. Note in particular the abrupt horizontal upper terminations of the cross-bedded facies in many places. The facies are identified as follows: (i) eolian dune cross-bedded sandstones; (ii) eolian sand sheet sandstones; (iii) paleosols; (iv) sandy interdune deposits (damp interdunes); (v) silty interdune deposits (wet interdunes). Modified from Melvin et al. (2005).

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in pronounced high angle (>30°), grain size–segregated crosslaminations. Those laminations may be very closely spaced (“pin-striped,” sensu Fryberger and Schenk, 1988) (Fig. 18A) and inversely graded, where they are interpreted as wind-ripple laminations (subcritically climbing translatent strata: Kocurek and Dott, 1981) that formed on the slip faces of eolian dunes. Alternatively, the cross-laminations are a few centimeters thick, in which case they are better interpreted as grain flow cross-strata (sensu Kocurek and Dott, 1981) that formed by gravity sliding of sand down the dune slip faces. These eolian dune cross-bedded deposits are readily identified on downhole image logs, which, crucially, allows this facies to be recognized in uncored wells. Eolian sand sheet deposits are also widely recognized within the upper Unayzah A member. They comprise fine- to

medium-grained sandstones, with very well rounded and frosted grains of sand that occur in low angle to flat, grain size–segregated laminations (Fig. 17, 90–98 ft; Fig. 18B). Many sets of these laminations have a pin-striped appearance, suggesting an origin from wind-ripple migration. Flat-based, lenticular (convex-up) accumulations of coarse-grained, well rounded sand are also observed within these deposits. It is possible in places to distinguish very low angle truncations separating different sets of the low-angle laminated sand. Some examples of this facies display disruption of the laminae, attributed to plant roots, as well as to other unknown forms of faunal bioturbation (insects? spiders?) (Fig. 18B). All of these characteristics compare favorably with features seen in eolian sand sheets, as they were described by Fryberger et al. (1979).

C A

B

D

Figure 18. Core photographs illustrating facies representative of the eolian erg center to erg margin facies association in the upper Unayzah A member. (A) Pronounced high-angle cross-lamination of the eolian dune facies. These laminations include very closely spaced (pin-striped) wind ripple laminations (wr) and thicker grain flow laminations (gf). (B) Very low-angle wind ripple lamination (wr) of the eolian sand sheet depositional facies. Note: low-angle truncations (arrowed) and subtle high-angle disruptions (insect burrows?) (arrowed, ib). These laminated sands overlie an interval at the bottom of the photograph that shows indistinct adhesion ripple (ar) structures. (C) Irregular crinkly laminations that are associated with damp interdune settings. White patches are irregular occurrences of anhydrite cement. (D) Disrupted and brecciated sandstone typical of many paleosols found within the arid eolian facies association of the upper Unayzah A member.

Late Paleozoic Gondwanan glaciation in Saudi Arabia Sandy interdune deposits (damp interdunes) are commonly recognized in the upper Unayzah A member. They consist of moderately sorted, fine- to medium-grained sandstones that contain a heterogeneous assemblage of depositional structures, including irregular to crinkly, variably continuous laminations and very thin (centimeter scale) beds with well-developed adhesion ripple laminations (Fig. 17, 84–90 ft; Fig. 18C). The occurrence of adhesion ripple laminations suggests that these sediments were laid down upon a damp substrate (cf. Kocurek and Fielder, 1982). The irregular nature of the laminations is reminiscent of textures observed in damp interdune, or sandy sabkha, environments (e.g., see Fryberger et al., 1983). Given the association of these sediments with facies (described above) that were clearly deposited in an arid, eolian-dominated setting, these irregularly laminated sandstones of the upper Unayzah A member are similarly interpreted to represent a damp interdune environment in close proximity to the paleo–water table. Silty interdune deposits (wet interdunes) comprise siltstones and silty, very fine–grained sandstones showing variably continuous to lenticular and crinkly lamination. In places small-scale ripple cross-lamination is seen, suggesting subaqueous deposition. These sediments occur in intervals that are rarely >1 ft (30 cm) to 2 ft (60 cm) thick. They show many depositional similarities to the semiarid lake deposits described above. The significant difference lies in the thickness of their occurrence. They are thus interpreted not as ephemeral (playa) lake sediments but rather as the deposits of temporary, very shallow interdune ponds that formed at times when the water table rose above the depositional surface within an otherwise arid (desert) environment. Sandy paleosol horizons have also been recognized at a number of localities within the eolian-dominated upper Unayzah A member. They are 0.3–4.0 ft (0.09–1.2 m) thick (Fig. 17, 53–57 ft, 126 ft, 151 ft), and comprise poorly sorted, very fine to medium-grained sandstones, with patchy carbonate cementation and showing varying degrees of disrupted texture (Fig. 18D). These paleosol horizons suggest that the paleo–water table was relatively high at the time of their formation. Individually, the above-described eolian and eolian-related facies occur commonly throughout the upper part of the Unayzah A member. The mutual associations of these facies, as well as their abundances relative to each other, vary markedly from well to well, and indeed from field to field across the study area. Thus in some places the upper part of the Unayzah A member is dominated (almost to the exclusion of any other facies) by stacked eolian dune cross-bedded sandstones that are clearly seen in both core and image-log data. These dune-dominated occurrences are interpreted as representative of eolian erg-center settings. Elsewhere a much greater mix of the various eolian facies (described above) is seen: those mixed facies associations are interpreted as eolian erg-margin deposits. Similar eolian-related facies associations have been described from the Permian Cedar Mesa Sandstone in Utah by Mountney and Jagger (2004). If well 21 (Fig. 17) is taken as a typical example of a well within which the facies of the upper Unayzah A member are eolian (as opposed to

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ephemeral fluvial) dominated, the repeatability of facies is clear, and the facies boundaries are seen to be sharp (Fig. 17). In particular, the eolian dune cross-bedded sandstones are in many cases abruptly truncated by a horizontal contact with the overlying facies (Fig. 19A; Fig. 17, 24 ft, 42 ft, 125 ft, 151 ft, 163 ft). These abrupt horizontal truncations of eolian cross-strata are identified as Stokes surfaces (after Stokes, 1968) and are attributed to a rising water table in the Permian eolian depositional environment (see Fryberger et al., 1988). The stratigraphic significance of the identification in the core of such rises in the paleo–water table was emphasized by Melvin et al. (2005). Subsidiary to the current work, a field-scale study was undertaken of the Unayzah A member at the southern end of the giant Ghawar structure in eastern Saudi Arabia that included wells 15 through 20 (see Fig. 1A). Details of that study are discussed by Melvin et al. (2010). There, the lower part of the Unayzah A member (described above as dominated by ephemeral lake facies) was seen in cores from many of the wells to terminate upward in a thin (decimeter scale), subtle upward-fining trend, with a passage from silty, very fine–grained sandstone to siltstone sensu stricto (Fig. 19B). That upward-fining trend can be interpreted as a deepening of the lower Unayzah A lake. The surface representing the top of this vertical, upward-fining trend among the wells has been interpreted further to represent in a lateral dimension the maximum extent of the lake (MEL) (Melvin et al., 2010). Above this MEL horizon, each well within the subsidiary study area displays its own assemblage of the various upper Unayzah A member eolian depositional facies (described above). Within these assemblages, respectively, a number of cyclical rises in the paleo–water table can be recognized based on the identification of Stokes surfaces in core, as well as on image logs, and the general associations of the various facies. Crucially, when the data from each well are hung stratigraphically using the MEL horizon as a datum, many of the interpreted rises in the paleo–water table that are recognized above the MEL in the various wells correlate exactly, irrespective of the depositional facies within which they are found (Fig. 20) (Melvin et al., 2010). This correlation extends along a distance of ~65 km. Thus, a correlatable layering scheme is crucially established that is related to paleo–water table fluctuations within the eolian-dominated Unayzah A member. The correlation demonstrably carries through different depositional facies tracts and is intrinsically established on sequence stratigraphic principles. Terminal Facies of the Uppermost Unayzah A Member The preceding discussion has described how the Unayzah A member is characterized by a wide variety of different depositional facies that generally fall into three distinctive facies associations, all of which are strongly indicative of deposition within a semiarid to arid setting. Thus this member almost universally in its lower part comprises playa lake deposits that pass upward into either ephemeral stream facies or eolian erg to erg margin facies. The ultimate termination of these conditions is represented in

Figure 19. Core photographs illustrating features of stratigraphic significance within the eolian facies association of the upper Unayzah A member. (A) Abrupt upper horizontal termination of pronounced eolian dune cross-bedding. This is a “Stokes surface” and is overlain in this example by sand sheet and interdune facies. It represents a rise through the dune deposits of the paleo–water table. Note: Large white spots are postdepositional nodules of anhydrite. (B) Upwardfining of very fine–grained sandstones into siltstone at the top of the lower Unayzah A member ephemeral lake deposits in well 18. The abrupt contact with overlying coarser (eolian) sandstones defines the maximum extent of the lake (MEL) in this well. See text for full discussion.

A

B

Figure 20. Stratigraphic cross section through the Unayzah A member in selected wells at the southern end of the Ghawar structure (see Fig. 1A). The section has the MEL (maximum extent of the lake) horizon (see text for discussion) as its datum. Note how a number of maximum wetting horizons above the MEL can be identified (from facies relationships seen in core and image logs), and that these appear to be correlatable along the length of the section, irrespective of the facies within which they are found. These wetting cycles are interpreted to be related to fluctuations in the paleo–water table. PKU—pre-Khuff unconformity.

Late Paleozoic Gondwanan glaciation in Saudi Arabia many wells across the study area by very different and distinctive facies at the top of the Unayzah A member (Fig. 21). Thus in well 24 the stratigraphically highest eolian deposits (Fig. 21A, 1–9 ft) are overlain by 21 ft (6.3 m) of finer-grained “interdune” deposits that differ from those normally seen in the upper Unayzah A member in that they display minor evidence of burrowing. Those sandstones are overlain by a highly distinctive interval 7 ft (2.1 m) thick, showing an abundance of burrows (Fig. 21A, 32–39 ft; Fig. 22A). Above this intensely bioturbated zone there occurs in well 24 a thick development of paleosols (Fig. 21A, 39–83 ft). Those paleosols appear to comprise at least five individual paleosol deposits each of which displays a number of probable soil “zones.” The facies characterizing these zones vary from very argillaceous, poorly sorted sandstones with an abundance of cutans and possible rooting features to extremely well developed, massive to nodular silcretes (e.g., Fig. 21A, 68–70 ft; Fig. 22B). This 44 ft (13.2 m)–thick interval of paleosols is superseded in this well by a heavily rubbleized interval of red silty mudstone that displays in its lower part an abundance of very thin (millimeter scale) laminations of fine- to mediumgrained sandstone. This sequence in well 24 can be compared favorably with the highest part of the upper Unayzah A member in well 21. In the latter well, some 300 km distant from well 24 (see Fig. 1A), the highest eolian dune sand is overlain by ~20 ft (6.0 m) of paleosol deposits (Fig. 21B, 10–30 ft). Those rocks subsequently pass upward through some upward-fining sandstones into an interval of low-angle laminated sandstone wherein the laminations are severely disrupted by intense burrowing (Fig. 21B, 39–45 ft; Fig. 22C). Overlying these burrowed sediments is another interval of thickly developed paleosols (Fig. 21B, 45–76 ft). Again, these paleosols appear to comprise a number (five or six) of stacked, individual paleosol deposits, within which various facies representing soil zonation occur. Those zones are similar to those identified in well 24 (see above) in that they vary from very argillaceous, poorly sorted sandstones with an abundance of cutans and apparent root-related structures to well-developed massive to nodular silcretes (e.g., Fig. 22D). It appears that, notwithstanding the great distance that separates wells 24 and 21, there are great similarities between the two wells in terms of the facies associations in the uppermost parts of the Unayzah A member. Specifically there appears to have been a thin, probable marine event, identified by intensive burrowing, that was followed in both cases by the development of thick, stacked paleosol deposits. In well 24 these paleosols are overlain by fine-grained red beds; in well 21 there is a 5 ft gap of non-recovery of core that is superseded by sandstones of shallow-marine origin and ascribed to the Basal Khuff Clastics member of the Khuff Formation (see later discussion).

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a large number of depositional facies are readily identified. It is also evident that most of those facies reflect deposition within a significantly arid environment. Thus, depending upon paleogeographical location within the study area, depositional facies associations reflect either (1) arid to semiarid conditions that are characteristic of terminal alluvial fans and bajadas or (2) widespread eolian erg systems, both of which encroached upon ephemeral lakes (playas) whose dimensions fluctuated throughout time. Stratigraphic analysis of one of these erg systems by Melvin et al. (2010) revealed a high degree of cyclicity among facies that results in a layering scheme that is correlatable both within and between depositional facies tracts (see earlier discussion). This internal stratigraphy within the eolian-dominated Unayzah A member has been interpreted to be a direct reflection of fluctuations in the level of the paleo–water table throughout the time of its deposition. These correlatable rises in the paleo–water table are intrinsically allostratigraphic in nature, and it is tempting to consider their cyclicity as possibly having its basis in the Earth’s orbital fluctuations. If the Gondwanan Permo-Carboniferous ice sheet was still significantly (albeit very distally) in existence during deposition of the Unayzah A member, and if such orbital fluctuations had an impact on the melting of the ice and the consequent postglacial global transgression, then it may be that even within the terrestrial rocks of the Unayzah A member there is a record of the pulsatory nature of that transgression. It appears to be highly significant in this regard that the stratigraphically highest of these terrestrial deposits in locations so widely separated as well 24 and 21 (Fig. 1A) are superseded by a thin zone of sandstone that displays a distinctly marine signature in the form of intense burrowing activity (Figs. 21 and 22). This would suggest that the episodic rises in base level inferred from the cyclicity identified from rises in the paleo–water table did indeed ultimately manifest themselves with a breakthrough of marine waters at the very end of the time of deposition of the Unayzah A member. It is clear that the end of deposition of the Unayzah A member in most areas was marked by a (probably prolonged) period of minimal deposition to nondeposition. This was reflected initially by the intensely bioturbated sandstone (indicating low rates of sedimentation) and was manifest ultimately by the development of the very thick and pervasive paleosols seen in many places at the top of this stratal unit. The widespread extent of those paleosols testifies to the development of a sustained, high water table that may indicate a change in climate (decreasing aridity). The significance of this inferred prolonged period of minimal deposition to nondeposition in terms of the overall tectono-stratigraphic evolution of Saudi Arabia is discussed in the final section of this paper. BASAL KHUFF CLASTICS MEMBER OF THE KHUFF FORMATION

Origin and Evolution of the Unayzah A Member It is clear from the foregoing discussion that the Unayzah A member is a highly complex stratigraphic unit, within which

Senalp and Al-Duaiji (1995) observed that the Unayzah Formation at outcrop is truncated by an unconformity known as the pre-Khuff unconformity. In the subsurface of the study

Figure 21. Core logs through the uppermost Unayzah A member at (A) well 24 and (B) well 21. These wells are ~300 km apart. Nonetheless, they display a remarkable similarity in their stratigraphic organization. Note in particular the occurrence in each well of a pronounced bioturbated sandstone zone (bs) ~25–30 ft (7.5–9.0 m) above the highest occurrence of eolian dune cross-bedded sandstone (dxb). In each case the bioturbated zone is overlain by a similar thickness (30–35 ft, 9.0–10.5 m) of paleosols (ps) that display similar internal stratigraphic organization. See text for full discussion.

Late Paleozoic Gondwanan glaciation in Saudi Arabia

A

B

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C

D

Figure 22. Core photographs showing aspects of the uppermost Unayzah A member in wells 24 and 21. (A) Intensely bioturbated sandstone in well 24 (Fig. 21A, 33–38 ft). (B) Pervasive development of silcrete in well 24 (Fig. 21A, 68–70 ft). (C) Intensely bioturbated sandstone in well 21 (Fig. 21B, 38–43 ft). (D) Pervasive development of silcrete in well 21 (Fig. 21B, 53 ft).

area (Fig. 1A), this unconformity separates the sandstones of the Unayzah from the overlying Khuff Formation. Although the Khuff Formation is widely recognized as a carbonate-dominated stratigraphic unit (Sharland et al., 2001; Vaslet et al., 2005), its lowermost deposits in many places are characterized by a series of alternating sandstones, shales, and thin carbonates that sit directly upon the pre-Khuff unconformity. Those deposits are described herein as the Basal Khuff Clastics member of the Khuff Formation. They include the unit identified by Senalp and Al-Duaiji (1995) and Vaslet et al. (2005) as the Ash Shiqqah member of the Khuff Formation. The depositional facies of the Basal Khuff Clastics member change in their character increasingly in a westward direction, as will be discussed below. Furthermore, the higher carbonate members of the Khuff Formation eventually overstep the Basal Khuff Clastics member, as well as the Unayzah and ultimately all older Paleozoic deposits, in the same direction. In extreme westerly locations, these upper Khuff carbonates rest directly upon Proterozoic rocks of the Arabian Shield.

Upward-Fining Units of the Basal Khuff Clastics In many places in the western part of the study area, and exemplified at wells 24 and 23 (Fig. 23A and Fig. 23B), the preKhuff unconformity is abruptly overlain by pebble conglomerates and very coarse–grained, pebbly sandstones that fine upward over several feet into argillaceous fine-grained sandstones and (ultimately) dark gray mudstones. The upward-fining sandstones and conglomerates have clasts that are variably subrounded to angular in nature and polymict in composition. Well 24, for example, displays an assemblage of angular detritus, including quartz, feldspar, jasper, granite, sandstone, siltstone, and mudstone. The mudstones that cap these upward-fining units in places display rooted textures in their upper part and commonly pass up into thin bioturbated sandstone and associated bioturbated mudrocks with calcareous concretions (e.g., well 23, Fig. 23B). The upward-fining character of these coarse-grained sediments in the western parts of the study area, and the lack of marine indicators except higher in the section, suggest that the

Figure 23. Selected core logs through the Basal Khuff Clastics member across the study area. (A) In well 24 the pre-Khuff unconformity (PKU) is overlain by a coarse breccia that fines upward into a highly carbonaceous paleosol. That in turn is overlain by ~5 ft (1.5 m) of thin-bedded marine bioturbated sandstones that pass up into gray mudstones (well 24). (B) In well 23 the pre-Khuff unconformity is abruptly overlain by upward-fining pebbly sandstone that passes upward into gray pedogenically altered mudrock. That soil is overlain by very thin, bioturbated fine-grained sandstone. (C) At well 15 the pre-Khuff unconformity is overlain by a mudrock-dominated interval that contains a few significant bioturbated sandstones. The mudrock facies show a cyclicity wherein fissile shales pass upward into more blocky and highly carbonaceous mudstones. (D) At well 17 the section directly above the preKhuff unconformity is characterized by a number of upward-fining cycles each of which passes into well-developed argillaceous and carbonaceous soils. These are superseded by a mudrock-dominated succession that contains bioturbated sandstones in its lower part, and within which can be seen a similar subtle cyclicity as at well 15. (E) The Basal Khuff Clastics member at well 21, in the eastern part of the study area, solely comprises shallow-marine depositional facies, including bioturbated sandstones and mudstones, and cross-bedded sandstones. See text for full discussion of these wells and their relationships.

Late Paleozoic Gondwanan glaciation in Saudi Arabia lowermost parts of the Basal Khuff Clastics member in those areas are characterized by fluvial deposition. The coarseness of the sandstones is indicative of a significant drop in base level (and/or uplift in the source area) pursuant to the creation of the pre-Khuff unconformity. Paleosols of the Basal Khuff Clastics In a number of the wells that have cored the Basal Khuff Clastics member, distinctive paleosols are recognized that are, significantly, quite different in character from those soils identified and earlier described from the top of the Unayzah A member. Thus in wells 24 (Fig. 23A, 35–40 ft) and 17 (Fig. 23D, 11–28 ft) fine- to very fine–grained, poorly sorted and very argillaceous and carbonaceous sandstones are identified low in the Basal Khuff Clastics section, overlying upward-fining coarser-grained sediment described above. The example from well 17 comprises four small-scale, upward-fining beds that display a high degree of vertical rooted texture throughout (Fig. 23D, 11–28 ft). These paleosols of the Basal Khuff Clastics member generally occur low in that stratal unit, directly subjacent to rocks displaying evidence of a marine influence in their deposition. This association leads to their interpretation as representing coastal marshlands in low-lying (estuarine) areas. As such they represent the first signs of the onset of the Khuff marine transgression. Marine Sandstone Deposits of the Basal Khuff Clastics In well 21 (Fig. 1A) the thick paleosols described above from the uppermost Unayzah A member (see Fig. 21B) are separated by ~6 ft (1.8 m) of non-recovery of core from a unit that is at least 21 ft (6.3 m) thick, and comprises sandstones and mudstones that are entirely different in character (Fig. 23E, 12–33 ft). Those sediments are assigned to the Basal Khuff Clastics member in this well. They comprise (1) interbedded dark gray, silty mudstones and thin (centimeter scale), heavily bioturbated fine-grained sandstones (Fig. 24A); (2) massive, intensively bioturbated sandstones up to 3 ft (0.9 m) thick; and (3) cross-bedded, fine-grained sandstones that occur in bedsets up to 5 ft (1.5 m) thick, with individual sets up to 0.5 ft (0.15 m) thick (Fig. 24B), and which only rarely display any bioturbation. In well 17 (Fig. 23D), the paleosols described above from the lowermost parts of the Basal Khuff Clastics member are overlain by a heavily bioturbated sandstone unit 6 ft (1.8 m) thick that contains large (cobble-sized) clasts of very fine–grained silica rock (Fig. 24C). Those clasts are probably derived from nearby nodular silcrete horizons within the uppermost Unayzah A paleosol (see earlier discussion). In well 15 the pre-Khuff unconformity is seen clearly to truncate eolian sandstones of the Unayzah A member (Fig. 23C). This unconformity is overlain by a very thin (1–2 cm) sandstone bed, above which the sequence is dominated by dark gray carbonaceous mudstones. Within those mudstones a very thin (0.3 ft, 0.09 m) bioturbated sandstone (Fig. 24D) occurs ~5 ft above the pre-Khuff unconformity in this well (Fig. 23C, 13–14 ft). In well

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24 (Fig. 23A) the paleosol described above from the Basal Khuff Clastics is abruptly overlain by ~5 ft of thin-bedded (decimeter scale) coarse-grained sandstones that are burrowed throughout. The depositional facies described above from well 21 were laid down in a shallow-marine environment wherein the thin sets of cross-bedded sandstone represent offshore marine bars, and the various bioturbated facies are indicative of slightly deeper (or otherwise protected) conditions that allowed a proliferation of infaunal activity. The bioturbated sandstone in well 17 was likewise deposited in a marine setting. It is probable that the large cobble-sized clasts of siliceous rock were eroded by storms from silcretes in the uppermost Unayzah paleosols in nearby locations. The coarse-grained (i.e., sand-sized) marine detritus observed in wells 21 and 17 is not recognized in well 15. There, the first definitive indications of marine deposition are seen in the thin bioturbated sandstone that occurs ~5 ft above the pre-Khuff unconformity (Fig. 23C, 13–14 ft; Fig. 24D). Significantly, palynological data from that zone show an influx of abundant scolecodont debris that can be interpreted as indicating a significant marine influx at that point (J. Filatoff, 2004, personal commun.). It is clear that a similar marine depositional signature is also evident in well 24 in the western part of the study area (Fig. 23A, 41–47 ft). Mudrocks of the Basal Khuff Clastics In wells 17 (Fig. 23D) and 15 (Fig. 23C) the marine sands described above are overlain by up to 40 ft (12 m) of dark gray mudrocks. They locally contain thin-walled bivalve shells and display a number of subtle cycles, each of which is ~5 ft (1.5 m) thick. Within each cycle, there is a passage upward from fissile, gray, silty mudstone to blocky and crumbly, highly carbonaceous and rooted mudstone (Fig. 23C, 15–47 ft; Fig. 23D, 45–77 ft). These fine-grained depositional cycles are interpreted to represent successive periods of lagoonal infill in a marginal marine setting, each one terminating in a coastal swamp. Origin and Evolution of the Basal Khuff Clastics Member of the Khuff Formation It is clear from the foregoing not only that the Basal Khuff Clastics member represents an extremely heterogeneous assemblage of depositional facies but also that that variation is geographically constrained. Thus in the southeast of the study area around well 21 the lowermost Basal Khuff Clastics member is specifically characterized by significant deposits of shallowmarine sandstones. Somewhat to the north and west, around well 17, the lowermost Basal Khuff Clastics consist of very different sediments, comprising relatively thin, upward-fining sandstones that have been interpreted above to be highly argillaceous and carbonaceous paleosols. Those paleosols are believed to have formed in a low-lying coastal plain or estuarine marsh environment. Above these paleosols in well 17, bioturbated marine sandstones are recognized, although they are not as fully developed as

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A

B D

C

Figure 24. Core photographs showing some characteristics of the marine sandstone facies of the Basal Khuff Clastics member. (A) Interbedded mudstones and fine-grained sandstones, displaying intense burrowing activity (well 21). (B) Cross-bedded sandstones of the offshore shallow-marine-bar environment (well 21). (See also Fig. 23E.) (C) Intensely bioturbated marine sandstone, containing a large clast of chert rock (Ch). The latter is considered to have been ripped out of nearby paleosol deposits by storms (well 17). (D) Thin, sharp-based bioturbated sandstone overlying gray mudrock (m). The sandstone is interpreted to be a relict shallow-marine storm deposit laid down below storm wave base (well 15). See text for further discussion.

those that occur in well 21; at well 15, shallow-marine sandstone is also recognized, but there it occurs only as a very thin bed (Fig. 24D) within an otherwise mudstone-dominated sequence. Those mudstones show a subtle cyclicity suggestive of repeated infill of lagoons in a coastal setting. In other words, the Basal Khuff Clastics in this location display a subtly less marine signature than equivalent deposits somewhat to the southeast at well 21. Even farther to the west in the study area, at well 23 the Basal Khuff Clastics member is represented in its lowermost part by

a fluvial deposit that is ultimately transgressed by a thin marine sandstone (Fig. 23B). At well 24 a basal fluvial breccia passes up into a paleosol that is in turn capped by ~5 ft (1.5 m) of thinbedded bioturbated sandstones (Fig. 23A, 32–47 ft). The clear distinction within the lowest parts of the Basal Khuff Clastics between shallow-marine depositional facies in the southeast and continental (fluvial) facies in the west demonstrates the gradual encroachment of the marine transgression from the southeast to the west in earliest Khuff times. The transgression

Late Paleozoic Gondwanan glaciation in Saudi Arabia reached its maximum expression with deposition of the various carbonate-evaporite cycles of the Khuff Formation, most recently described by Vaslet et al. (2005).

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tiate the much closer association of the Unayzah B member with the unnamed middle Unayzah member in mid- to high southerly latitudes, from the near tropical setting of the Unayzah A member. These observations are discussed further below.

PALEOMAGNETIC STUDIES DISCUSSION In the course of conducting these stratigraphic studies of the Unayzah and lowermost Khuff Formations, an opportunity arose to carry out a paleomagnetic pilot study on core samples from two of the wells, wells 7 and 22. A total of 38 samples were collected under appropriately controlled, nonmagnetic conditions, and analyzed with a view to determining whether or not the paleolatitude at the time of their deposition could be identified. The samples were taken from rocks that were independently assigned to the Unayzah B member, the un-named middle Unayzah member, and the Unayzah A member. The results of this pilot study are presented in Figure 25. Although the samples from each stratigraphic unit show a range of values, it is clear that each unit has its own signature with respect to its interpreted paleolatitude and that there is no overlap in the respective data sets. Thus the samples from the Unayzah B member were deposited in paleolatitudes represented by a mean value ~75° S. The unnamed middle Unayzah member appears to have been laid down when the study area was at a (mean) paleolatitude of ~55° S. Deposition of the Unayzah A member did not occur until the area of investigation lay in paleolatitudes of ~27° S. These data clearly reflect the northward drift of the Arabian plate throughout the Early to Middle Permian, as was independently documented by previous authors (Beydoun, 1991; Al-Fares et al., 1998). They also appear to clearly differen-

Unayzah A

Younger

Unnamed middle Unayzah member

Unayzah B

0

30 60 Paleolatitude (degrees S)

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Figure 25. Diagram based on paleomagnetic data, showing relative position with respect to paleolatitude of the Unayzah B member, the unnamed middle Unayzah member, and the Unayzah A member. These data sets individually display varying degrees of spread but nonetheless appear to show no overlap with each other. They illustrate clearly the northward drift of the Arabian plate from Asselian–earliest Sakmarian time (Unayzah B member) to Artinskian–Kungurian(?) times (Unayzah A member). The results were obtained from core plugs collected under controlled (nonmagnetic) conditions and were analyzed by E. Hailwood (CoreMagnetics).

Sharland et al. (2001) discussed how Tectonostratigraphic Megasequence TMS AP5 is the last megasequence of the Arabian plate to have been dominated by siliciclastic sediments (Unayzah Formation). Only the Basal Khuff Clastics member, overlying the pre-Khuff unconformity at the base of TMS AP6, represents any further siliciclastic deposition in Saudi Arabia prior to the northward drift of the Arabian plate into subtropical latitudes where carbonate and evaporite deposition predominated (cf. Beydoun, 1991; Al-Fares et al., 1998). TMS AP5 is marked at its base in Saudi Arabia by the Hercynian unconformity, which represents erosion associated with the mid-Carboniferous “Hercynian” tectonic inversion event (Al-Husseini, 2004). Those erosional processes were intensified with the almost coincident inception of the late Paleozoic Gondwanan glaciation in Arabia. The upper boundary of TMS AP5 is the pre-Khuff unconformity, dated as mid-Tatarian in Saudi Arabia by Stephenson and Filatoff (2000b). It represents erosion following the early Late Permian rifting associated with the creation of the Neotethys Ocean (cf. Bishop, 1995; Loosveld et al., 1996). It can thus be considered to be the “Break-up Unconformity” associated with the continental rifting and spreading of the Sanandaj-Sirjan and central Iran terranes from the Arabian plate (Sharland et al., 2001). Angiolini et al. (2003), however, believe that in the Sakmarian of Oman there are two tectono-eustatic transgressive events that are related to the end of the Gondwanan glaciation and the concomitant tectonic evolution during rifting and initial opening of the Neotethys Ocean. Those authors thus consider the “Break-up Unconformity” to be represented by an unconformity that is seen within TMS AP5, and which formed much earlier than the pre-Khuff unconformity (Angiolini et al., 2003). Clearly, the evolution of Tectonostratigraphic Megasequences TMS AP5 and TMS AP6 on the Arabian plate represents a complex interplay of tectonic, climatic, and eustatic relationships. An attempt is made below to discuss that evolution, in the context of the depositional and stratigraphic development of the various units described above within the Unayzah, as well as the Basal Khuff Clastics member of the Khuff Formation of Saudi Arabia. Inevitably, that discussion necessitates consideration of the equivalent stratal units elsewhere on the Arabian plate, specifically Oman. As has been discussed above, the late Paleozoic South Polar glaciation that affected so much of the Gondwanan continent is considered by Al-Husseini (2004) possibly to have commenced in Arabia as early as middle Carboniferous (late Visean–early Namurian or Serpukhovian) times, and to have persisted for 30–45 m.y. until the Early Permian (Sakmarian). This provided ample time for an unknown number of glacial advances, retreats,

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and readvances to have taken place, both in Saudi Arabia and Oman. Such multiple events are implicitly recognized in Oman from palynological evidence within the lower Al Khlata (Osterloff et al., 2004a), which presumably reflects different interstadial depositional periods. In Saudi Arabia the palynology is not so forthcoming. It is tempting to think that the glacial readvances that have been postulated herein in relation to the shear zones within the Unayzah C member may be correlative by some means with the Oman section. This is clearly an important area for future study. The final retreat phase of the Gondwanan glaciation in Arabia (and by implication the onset of climatic amelioration) is manifest in the great variety of predominantly glaciolacustrine depositional facies recognized and described above in the Unayzah B member in Saudi Arabia as well as in the upper Al Khlata Formation (including the Rahab Shale) of Oman (Osterloff et al., 2004a). Paleomagnetic evidence from the current study appears to suggest that at this time Saudi Arabia lay in high southerly latitudes ~75° S. Notwithstanding minor readvances such as have been recognized in the local occurrence of stratabound push moraine deposits (Fig. 7), the overall character of the Unayzah B member across most of the study area illustrates ongoing glacial retreat, evident in the continuing filling and spilling over of an environment that was dominated by glacial lakes. This wholesale flooding of the postglacial landscape in Saudi Arabia is reflected in the clearly “transgressive” (deepening) nature of the glaciolacustrine depositional facies associations. In the western part of the study area the Unayzah B is characterized predominantly by glacially induced deformation features. These are considered to reflect the longer term presence of significant bodies of ice in these locations, and possibly even an element of long-lived alpine glaciation associated with highlands of the Al Batin Arch. The uppermost deposits of the Unayzah B member across most of the subsurface of eastern Saudi Arabia (and of the Rahab Shale in Oman, Osterloff et al., 2004a) provide clear sedimentological evidence of the maximum melt-out of the Gondwanan ice sheet. It remains to be seen whether or not these rocks represent the maximum flooding event of the entire Tectonostratigraphic Megasequence TMS AP5. That in turn demands a consideration of the extent to which the postglacial transgression is reflected by purely climatic (i.e., warming) controls, or by overriding plate tectonic factors. The only maximum flooding surface (MFS) that has hitherto been formally recognized in TMS AP5 was identified in Oman by Sharland et al. (2001) as MFS P10, which is represented by a bioturbated shale directly below the Haushi Limestone. This is the “maximum flooding shale” of Guit et al. (1995) and occurs within the postglacial lower Gharif member (see Fig. 2). Stephenson and Osterloff (2002) showed that MFS P10 is marked by the occurrence of the acritarch Ulanisphaeridium omanensis, and Stephenson et al. (2003) subsequently assigned it to the OSPZ3b Subbiozone. Angiolini et al. (2006) showed that the marine deposits of the overlying Haushi Limestone are Early Permian (Sakmarian) in age and correlative with the OSPZ3c Subbiozone

of Stephenson et al. (2003). Melvin and Sprague (2006) proposed that the unnamed middle Unayzah member red beds in Saudi Arabia can be correlated with the lower Gharif member in Oman and therefore may be considered generally correlative with the OSPZ3 Zone. This implied that in Saudi Arabia MFS P10 would somehow have equivalence within the wholly terrestrial deposits of the unnamed middle Unayzah member of Melvin and Sprague (2006). Ongoing palynological studies are providing encouraging results that support this original proposition of Melvin and Sprague (2006) that there is at least some stratigraphic equivalence between the unnamed middle Unayzah member and the lower Gharif member (N.P. Hooker, 2008, personal commun.). In Oman, Angiolini et al. (2003) considered the paleoclimatic and paleotectonic characteristics of the Saiwan Formation at outcrop (equivalent in the subsurface to the Haushi Limestone at the top of the lower Gharif member). They concluded from paleontological, paleoecological, and petrographic evidence that the opening of the Neotethys Ocean commenced north of Oman in about midSakmarian times (Angiolini et al., 2003). From the foregoing discussion of the sedimentary history of the Unayzah Formation, it would appear nonetheless that these proposed early stages of the Neotethys marine transgression did not find recognizable expression in the terrestrial unnamed middle Unayzah member in Saudi Arabia. The significant drainage event postulated earlier as marking the end of Unayzah B member times, and, most significantly, the end of the Gondwanan glaciation in Saudi Arabia, may well have been initiated by distant tectonic events related to an incipient breach as a first step toward opening of the Neotethys Ocean. In Oman, Guit et al. (1995) suggested the presence of a regional intra-Gharif unconformity related to a drop in relative sea level at the top of the lower Gharif member, and Osterloff et al. (2004b) also described how the base of the middle Gharif marks an overall regression. In Saudi Arabia the contact between the unnamed middle Unayzah member and the overlying Unayzah A member is considered to be a sequence boundary that is at least disconformable on the subcropping formation (Melvin and Heine, 2004). It heralds the onset of a major shift in paleoclimatic conditions, under which deposition of the Unayzah A member became dominated by arid to semiarid eolian and ephemeral stream sedimentation. That climatic change was probably brought about by the northward drift of the Arabian plate into lower southerly latitudes (see Beydoun, 1991; Al-Fares et al., 1998). Northward drift is clearly reflected in the paleomagnetic data presented in the current paper (Fig. 25). Those data demonstrate a large amount of northward migration toward tropical latitudes and can be inferred to represent a significantly long period of time between deposition of the unnamed middle Unayzah member and the Unayzah A member. This in turn would suggest that the stratigraphic boundary between these two members represents a hiatus of considerable significance. The earlier discussion of the Unayzah A member has highlighted the identification in places of a number of zones within this stratigraphic unit, each of which is related to a demonstrable rise in the paleo–water table and is correlatable over tens of

Late Paleozoic Gondwanan glaciation in Saudi Arabia kilometers, irrespective of the depositional facies within which it occurs. The possibility has already been mooted that this cyclical rise in the paleo–water table through the sediments of the Unayzah A member may be related to a pulsed rise in distant sea level, and as such may be a phantom expression of ongoing marine transgression. The recognition of an apparently widespread bioturbated (and therefore probably marine) sandstone zone close to the top of the Unayzah A member is highly relevant in this regard. The middle Gharif member in Oman is considered the probable stratigraphic equivalent of the Unayzah A member in Saudi Arabia (Stephenson et al., 2003). It has been described by Osterloff et al. (2004b) as consisting “largely of non-marine clastics, indicated by red continental paleosols and non-marine facies.” Significantly, the latter authors also note that a “rapid vertical interdigitating nature of marine to non-marine clastics” has been recorded in the middle Gharif member from cores from central Oman. It is possible that such alternations of facies in Oman reflect a more marineward expression of the cyclicity that has been seen and described herein from the predominantly continental sediments of the Unayzah A. That cyclicity culminates with the prominent bioturbated (and therefore probably marine) zone that occurs in some wells in the uppermost parts of the Unayzah A member (cf. Fig. 21), suggesting final breakthrough in latest Unayzah A time by marine waters into the study area of eastern and central Saudi Arabia. Thereafter the uppermost Unayzah A member displays widespread development of generally thick paleosols, suggesting a prolonged period of nondeposition, with no marine indicators. This is interpreted to be a reflection of the up-doming that has been postulated by Sharland et al. (2001) in relation to the thermal uplift that preceded continental rifting and spreading, and which culminated in the regional pre-Khuff unconformity. It was noted earlier that the palynological changeover between OSPZ4 and OSPZ5 is “probably the greatest recorded in the Permian palynological succession in Oman and Saudi Arabia” (Stephenson et al., 2003). This palynological change finds expression in Oman also in changes in depositional (environmental), chemostratigraphic, and mineralogical (heavy minerals) signatures between the middle Gharif and upper Gharif members (Stephenson et al., 2003; Osterloff et al., 2004b). It is thus strongly implicit that a significant depositional hiatus exists between these two members. Osterloff et al. (2004b) noted how the upper Gharif member consists of four cycles (namely cycles 5–8), the uppermost of which displays evidence of “tidal/ estuarine” environments. The preceding discussion has demonstrated a similar change in Saudi Arabia from the widespread arid continental setting in the Unayzah A member to the mixed fluvial to shallow-marine setting of the Basal Khuff Clastics member. Furthermore, the Basal Khuff Clastics member exhibits a distinctive and different heavy mineral assemblage relative to the underlying sandstones of the Unayzah A (R.W.O’B. Knox, 2003, written commun.). It is thus evident that a significant depositional changeover exists between the Unayzah A member and the Basal

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Khuff Clastics member: the boundary for that changeover is manifest as the pre-Khuff unconformity. Osterloff et al. (2004b, p. 115) postulated some equivalence in the sequence development of the uppermost cycle of the upper Gharif member in Oman compared with the upper Basal Khuff Clastics member, while stressing that no age equivalence is implied. Indeed, Stephenson (2006) recently clarified how in Oman the upper Gharif member is assigned to palynozone OSPZ5 of Stephenson et al. (2003), whereas the Basal Khuff Clastics member in Saudi Arabia is assigned definitively to the younger palynozone OSPZ6 (see Fig. 2). It has been shown earlier in this paper how in Saudi Arabia the lowermost part of the Basal Khuff Clastics member displays very different depositional facies from place to place, ranging from shallow marine in the southeastern part of the study area to fluvial deposits in more westerly locations (Fig. 23). This geographical demarcation of facies clearly reflects the onset of fully marine transgression to the south and east of the study area. That transgression progressed steadily westward, as is reflected in (1) the diachroneity of deposits from the upper Gharif member to the Basal Khuff Clastics, (2) the heterogeneity of the facies tracts within the Basal Khuff Clastics, and, ultimately, (3) by the superposition of Khuff carbonates over Proterozoic basement rocks in western central Saudi Arabia, as discussed previously. The pre-Khuff unconformity marks the top of the Tectonostratigraphic Megasequence TMS AP5 of Sharland et al. (2001). Those authors describe how “progressive thermal uplift (the precursor to continental rifting and spreading) is interpreted to have occurred … culminating in the regional ‘pre-Khuff unconformity’ at the top of this megasequence” (Sharland et al., 2001, p. 85). The change in Oman from continental or marginal-marine deposits of the Gharif Formation to fully marine Khuff Formation carbonates in the Wordian (Broutin et al., 1995) (equivalent to early Kazanian) is interpreted “to reflect marine flooding following thermal collapse of the new Arabian plate passive margin with Neotethys” (Sharland et al., 2001, p. 89). Those authors observe that “in north Arabia the transition probably occurs slightly later, at the base of the Tatarian (equivalent to middle Capitanian) . . . possibly indicating a very rapid ‘unzipping’ effect from southeast to northwest.” The present paper presents data that fully support these concepts. Furthermore, it seems logical that in lithostratigraphic terms the upper Gharif in Oman is indeed an older equivalent of the Basal Khuff Clastics member in Saudi Arabia and marks the onset in Oman of the major transgression that ultimately took place across Arabia, pursuant to the creation of the Neotethys Ocean. CONCLUSIONS 1. In middle Carboniferous times the Arabian plate was subjected to an episode of major tectonic activity that was followed by a period of erosion and the creation of the Hercynian unconformity. These events were closely followed by the inception of the late Paleozoic Gondwanan glaciation in Arabia, which

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resulted in deposition of the Unayzah Formation upon the Hercynian unconformity. 2. The Unayzah C member, where present, sits directly upon the Hercynian unconformity. It varies greatly in thickness and comprises quartzose sandstones that were laid down during multiple retreat phases of the ice sheets upon glaciofluvial outwash braided plains. During subsequent intervening readvances of the ice sheets, these outwash deposits were overridden, cannibalized, and pushed into stacked piles of glacially tectonized (push moraine) deposits, manifest as relatively undeformed sandstones separated by distinct shear zones. The top of the Unayzah C member is an unconformity and represents the final subglacial surface of the Gondwanan ice sheet in Saudi Arabia. 3. The terminal melt-out phase of the Gondwanan glaciation is identified in the various facies of the Unayzah B member, which was deposited in high southerly latitudes, ~75° S, based on paleomagnetic data. These Unayzah B sediments comprise a number of different ice-proximal to ice-distal glaciolacustrine depositional facies. Facies associations consistently demonstrate progressive retreat of the ice and a concomitant sustained increase in flooding by meltwaters of an environment that was characterized by an abundance of glacial lakes. The top of the Unayzah B member can be taken to represent the climatic maximum flooding surface related to the melting of the Gondwanan ice sheets in Saudi Arabia and is considered equivalent to the top of the Rahab Shale in Oman. There is evidence to suggest a possible center of relatively prolonged upland glaciation in the western part of the study area, associated with the Al Batin Arch. 4. The unnamed middle Unayzah member is separated from the underlying Unayzah B member by a sharp stratigraphic boundary, considered to be a regional disconformity, which is interpreted to represent a significant drainage event that marked the end of the glaciation in Saudi Arabia. Paleomagnetic data suggest that the unnamed middle Unayzah member was laid down in paleolatitudes ~55° S. This member comprises red-brown, alluvial floodplain sandy siltstones, within which occur various isolated sandstone facies representing fluvial and possibly coldclimate eolian deposits. It is commonly capped by a paleosol horizon that represents a terrestrial highstand deposit, believed to be equivalent to the marine Haushi Limestone highstand deposit at the top of the lower Gharif member in Oman. That those paleosols represent a prolonged period of limited deposition or nondeposition is supported by the paleomagnetic evidence. The paleosols are therefore considered to mark a significant disconformity between the unnamed middle Unayzah member and the overlying Unayzah A member. 5. The lower Unayzah A member was deposited in widespread ephemeral (playa) lakes, whereas the upper Unayzah A sediments were deposited predominantly within arid to semiarid eolian dune fields and associated interdune deposits as well as within ephemeral stream systems and minor playa lakes. In places the eolian deposits display an internal stratigraphy that is related to cyclical fluctuations in the paleo–water table, and which is correlatable within and between significant facies tracts in this

stratal unit. Those rises in the paleo–water table possibly reflect ongoing pulsed phases of a distant marine transgression, an idea that is given support by the recognition of a heavily bioturbated sandstone zone at the very top of the Unayzah A member in two wells at either end of the study area. That zone is thought to represent final breakthrough of the transgressing marine waters. In most places the very highest parts of the Unayzah A display thick, well-developed paleosol horizons. These represent an extended period of minimal deposition and are considered to reflect the long period of uplift that was related to thermal doming prior to rifting and collapse that led to the creation of the pre-Khuff unconformity. The Unayzah A member is considered equivalent to the middle Gharif member of Oman. 6. The Unayzah A member is truncated by the pre-Khuff unconformity, which marks the upper boundary of Tectonostratigraphic Megasequence TMS AP5 in Arabia. It is overlain by sandstones and shales of the Basal Khuff Clastics member of the lowermost Khuff Formation. The depositional facies of the Basal Khuff Clastics are characterized in southeastern areas by shallowmarine sandstones and shales that pass westward into more terrestrial facies (fluvial), reflecting the onset of a major westwarddirected marine transgression. That transgression reached its fullest development with the widespread marine carbonates of the Khuff Formation. It is an expression of the successful opening of the Neotethys Ocean, and as such represents a tectonically related major flooding event. 7. The late Paleozoic Gondwanan glaciation per se in Saudi Arabia is represented in part by the Hercynian unconformity, as well as by the deposition and deformation of the Unayzah C member. Its final retreat from Saudi Arabia is seen in the rocks of the Unayzah B member, the top of which represents the maximum glacial melt-out. This is the maximum expression of postglacial transgression that can be ascribed solely to climatic (warming) factors. The end of the glaciation in Saudi Arabia is marked by an inferred dramatic drainage event, manifest in the disconformity that separates the glaciolacustrine Unayzah B member from the terrestrial deposits of the unnamed middle Unayzah member. Thereafter, the rocks are continental in nature, and their relationships with distant marine transgression (for which there is tantalizing evidence) must necessarily incorporate the likelihood of significant tectonic influences. These tectonic influences are ultimately manifest in the pre-Khuff unconformity. Above that unconformity the rocks of the Basal Khuff Clastics member of the Khuff Formation reflect the marine transgressive flooding associated with the opening of the Neotethys Ocean. ACKNOWLEDGMENTS We acknowledge the Saudi Arabian Ministry of Petroleum and Mineral Resources and the Saudi Arabian Oil Company (Saudi Aramco) for granting permission to publish this paper. The evolution of our thoughts on the Unayzah Formation has benefited from many discussions with many colleagues at Saudi Aramco. In particular we wish to acknowledge Mark Prudden,

Appendix: Core Log Legend Planar lamination

Mud cracks

Cross lamination Burrows Trough crosslamination Rooted horizons Low-angle lamination Current ripples

Pedogenic “slickoliths” Floating sand grains

Climbing ripples

Mud clasts

Coarse sand lamination

Pebbles, cobbles

Wispy lamination

Chert

Adhesion ripple

Carbonaceous layer

Graded bedding

Upward-shoaling interval

Dish structure with elutriation pillars Soft sediment deformation

Stylolites

Flame structures

Vertical fractures

Soft sediment loading

Low-angle shear

High-angle stylolites

Grain size G S Z M Gravel Sand Silt Mud/clay Figure A1. This figure provides a legend to which reference should be made for all figures throughout the paper comprising sedimentological core logs.

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Kent Norton, and Roger Price of Saudi Aramco’s Exploration Organization for much stimulating debate. John Filatoff, Nigel Hooker, and Merrell Miller provided biostratigraphic information, and Stephen Franks shared petrographic data regarding these enigmatic rocks. Peter Sharland also provided valuable insights on the occurrence of late-glacial drainage events during the Pleistocene glaciation. Ernie Hailwood of CoreMagnetics provided the analyses of the samples selected for the paleomagnetic (paleolatitude) studies that are referenced herein. We nonetheless accept sole responsibility for the ideas presented in this work. We thank George Grover and the Saudi Aramco Publications Review Committee for their time and effort, and for insightful comments in reviewing this paper. Kathleen Haughney of the Saudi Aramco Exploration and Producing Information Center was her usual cheerful self in chasing down some of the more elusive reference material. Gene Cousart of Aramco’s Cartography Department demonstrated outstanding commitment and professionalism in his approach to producing the figures that are presented in this paper. Ali Al-Zahrani and Hadi Al-Uraij are thanked for their tireless efforts in arranging for the layout of many thousands of feet of Unayzah core in the Saudi Aramco Core Storage facility. REFERENCES CITED Ahlbrandt, T.S., and Andrews, S., 1978, Distinctive sedimentary features of cold-climate eolian deposits, North Park, Colorado: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 25, p. 327–351, doi: 10.1016/0031 -0182(78)90048-2. Aitken, J.F., Clark, N.D., Osterloff, P.L., Penney, R.A., and Mohiuddin, U., 2004, Regional core-based sedimentological review of the glacially influenced Permo-Carboniferous Al-Khlata Formation, South Oman Salt Basin, Oman: GeoArabia, Abstract, v. 9, no. 1, p. 16. Al-Belushi, J.D., Glennie, K.W., and Williams, B.P.J., 1996, Permo-Carboniferous glaciogenic Al Khlata Formation, Oman: A new hypothesis for origin of its glaciation: GeoArabia, v. 1, p. 389–404. Al-Fares, A.A., Bouman, M., and Jeans, P., 1998, A new look at the middle to Lower Cretaceous stratigraphy, offshore Kuwait: GeoArabia, v. 3, p. 543–560. Al-Hajri, S.A., and Owens, B., eds., 2000, Stratigraphic Palynology of the Palaeozoic of Saudi Arabia: GeoArabia Special Publication 1, Gulf PetroLink, Bahrain, 231 p. Al-Husseini, M.I., 2004, Pre-Unayzah unconformity, Saudi Arabia, in AlHusseini, M.I., ed., Carboniferous, Permian and Early Triassic Arabian Stratigraphy: GeoArabia Special Publication 3, Gulf PetroLink, Bahrain, p. 15–59. Angiolini, L., Balini, M.E., Garzanti, E., Nicora, A., and Tintori, A., 2003, Gondwanan deglaciation and opening of Neotethys: The Al-Khlata and Saiwan Formations of Interior Oman: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 196, p. 99–123, doi: 10.1016/S0031-0182(03)00315-8. Angiolini, L., Stephenson, M.H., and Leven, E.J., 2006, Correlation of the Lower Permian surface Saiwan Formation and subsurface Haushi limestone, Central Oman: GeoArabia, v. 11, p. 17–38. Bates, R.L., and Jackson, J.A., 1980, Glossary of Geology (2nd edition): Falls Church, Virginia, American Geological Institute, 751 p. Bennett, M.R., 2001, The morphology, structural evolution and significance of push moraines: Earth-Science Reviews, v. 53, p. 197–236, doi: 10.1016/ S0012-8252(00)00039-8. Beydoun, Z.R., 1991, Arabian Plate Hydrocarbon Geology and Potential—A Plate Tectonic Approach: American Association of Petroleum Geologists Studies in Geology 33, 77 p. Bishop, R.S., 1995, Maturation history of the Lower Palaeozoic of the Eastern Arabian Platform, in Al-Husseini, M.I., ed., Middle East Petroleum Geosciences GEO94: Gulf PetroLink, Bahrain, v. 1, p. 180–189.

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MANUSCRIPT ACCEPTED BY THE SOCIETY 21 DECEMBER 2009

Printed in the USA

The Geological Society of America Special Paper 468 2010

Environmental and paleogeographic implications of glaciotectonic deformation of glaciomarine deposits within Permian strata of the Metschel Tillite, southern Victoria Land, Antarctica John L. Isbell* Department of Geosciences, University of Wisconsin–Milwaukee, Milwaukee, Wisconsin 53021, USA

ABSTRACT Popular reconstructions of late Paleozoic glaciation depict a single massive ice sheet centered over Victoria Land and extending outward over much of Gondwana. This view is untenable, as interpretations presented here indicate that glaciogenic strata in southern Victoria Land were deposited in a glaciomarine setting, and that ice entered the area from at least two different ice centers on opposite sides of the depositional basin. Reports from other areas also reveal that multiple ice sheets, ice caps, and alpine glaciers diachronously waxed and waned across Gondwana during the Carboniferous and Permian. Glaciogenic rocks of the Lower Permian Metschel Tillite contain glaciotectonic structures and glaciogenic deposits that include (1) sheared diamictites, (2) thrust sheets, (3) massive and stratified diamictites, and (4) sheet sandstones. These features formed as subglacial deforming beds, thrust moraines at glacial termini, and as glaciomarine deposits associated with temperate glaciers. A glaciomarine setting, rather than a glaciolacustrine setting, is suggested, owing to the abundance of meltwater plume deposits. A wedge-shaped sandstone body at the base of the overlying Weller Coal Measures was deposited as a grounding-line fan. Results of this study imply deposition in ice-marginal glaciomarine settings from ice radiating out of multiple glacial centers. These findings are significant because multiple glaciers, covering a given area, contain considerably less ice volume than a single massive ice sheet. Therefore, the waxing and waning of multiple ice masses during the late Paleozoic would have influenced global climate and eustatic sea level much differently than would have a single massive Gondwanan ice sheet. INTRODUCTION

Permian (Fig. 1; Veevers, 1994, 2001; Zeigler et al., 1997; Hyde et al., 1999; Scotese et al., 1999). In these models, ice flowed radially outward from a glacial spreading center over Victoria Land, Antarctica, and extended, in one direction, to glaciomarine margins in the Ellsworth Mountains and South Africa (Lindsay, 1970; Barrett, 1991; Veevers, 2001). In the opposite direction, models show ice flowing out of northern Victoria Land to terrestrial ice

Many paleogeographic reconstructions of the late Paleozoic Gondwanan Ice Age (LPGIA) depict an immense ice sheet covering Gondwana during the Mississippian, Pennsylvanian, and *[email protected]

Isbell, J.L., 2010, Environmental and paleogeographic implications of glaciotectonic deformation of glaciomarine deposits within Permian strata of the Metschel Tillite, southern Victoria Land, Antarctica, in López-Gamundí, O.R., and Buatois, L.A., eds., Late Paleozoic Glacial Events and Postglacial Transgressions in Gondwana: Geological Society of America Special Paper 468, p. 81–100, doi: 10.1130/2010.2468(03). For permission to copy, contact [email protected]. ©2010 The Geological Society of America. All rights reserved.

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Africa

India Au

S. America

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str ali a

Antarctica VL

CTM EM

Figure 1. (A) Map showing the distribution of glacial basins in Gondwana. (B) Gondwana reconstruction after Powell and Li (1994), showing the location of a hypothetical ice sheet covering Gondwana (modified from Ziegler et al., 1997) with an ice spreading center over Victoria Land, Antarctica (Lindsay, 1970; Barrett, 1991). Hypothetical flow directions are from models proposed by Lindsay (1970), Crowell and Frakes (1971), Barrett (1991), and Veevers (2001). CTM—central Transantarctic Mountains, EM—Ellsworth Mountains, NZ—New Zealand, VL—Victoria Land.

NZ

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km 1500

Antarctica

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NZ Gondwanan Ice Sheet

margins in southern Australia and glaciomarine margins in Tasmania (Fig. 1; Crowell and Frakes, 1971; Barrett, 1991; Veevers, 2001). Throughout Antarctica, early workers identified upper Paleozoic diamictites as tillites deposited subglacially from the terrestrial portion of the ice sheet as it waxed and waned across Gondwana (e.g., Lindsay, 1970; Barrett and Kyle, 1975; Barrett et al., 1986). The occurrence of an ice spreading center in Victoria Land supplying ice to the massive Gondwanan Ice Sheet is central to many paleogeographic and climatic models for the late Paleozoic. These models are based on work conducted prior to the advent of modern glacial facies models, and therefore their validity has not been rigorously tested in Victoria Land, the proposed terrestrial center of the ice sheet. In Antarctica (Fig. 2), upper Paleozoic glacial deposits thin from the central Transantarctic Mountains into southern Victoria Land. There, glaciogenic deposits are discontinuous (Barrett and Kyle, 1975; Collinson et al., 1994). In southern Victoria Land, thick glaciogenic deposits are reported as glacial terrestrial valley fills (Barrett, 1972; Barrett and Kyle, 1975; McKelvey et al., 1977; Barrett and McKelvey, 1981). Barrett and McKelvey (1981) suggested that the valleys preferentially preserved the deposits from erosion during postglacial rebound, whereas, outside the valleys, partial if not complete erosion of the glacial

Victoria Land Ice Spreading Center 0

km 1500

deposits occurred prior to deposition of the overlying Lower Permian Weller Coal Measures. The occurrence of the glacial valleys and regional thinning of the glacial strata toward Victoria Land led Lindsay (1970) and Barrett (1991) to conclude that southern Victoria Land lay adjacent to a late Paleozoic glacial spreading center and that deposition occurred from the terrestrial portion of the ice sheet (Fig. 1B). However, data presented herein suggest that some of the relief on the preglacial contact separating upper Paleozoic glacial rocks above from Devonian strata below resulted from glaciotectonic thrusting along concave upward shear planes rather than from erosion and development of glacial valleys, and that deposition occurred within a glaciomarine setting. Even though glaciotectonic structures and their significance have been recognized in Gondwanan rocks in South Africa and South America (e.g., Visser and Loock, 1982; Rocha-Campos et al., 2000), descriptions of pre-Pleistocene glaciotectonic features are scarce. Deformed upper Paleozoic glaciogenic rocks also occur in Antarctica. The identification and interpretation of these strata are enigmatic, as deformation has been attributed to overriding of the sediments by Late Carboniferous or Permian ice, slumping along glacial valley walls, and/or owing to tectonic upheaval (Barrett, 1972; Barrett and Kyle, 1975; McKelvey

Permian deposits, southern Victoria Land, Antarctica

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B

Figure 2. (A) Map of Antarctic sites showing the location of southern Victoria Land. (B) Location map of southern Victoria Land showing Kennar Valley, Mount Metschel, and Mount Ritchie.

et al., 1977; Barrett and McKelvey, 1981; McElroy and Rose, 1987; Spörli, 1992). This paper presents sedimentologic data collected from Mount Metschel, Mount Ritchie, and Kennar Valley in southern Victoria Land (Fig. 2) during the 2000–2001 and 2001–2002 austral field seasons. Here, strata of the Devonian Aztec Siltstone, the Permian Metschel Tillite, and the basal few meters of the Weller Coal Measures (Fig. 3) are described and interpreted (1) to determine whether a major late Paleozoic ice sheet spreading center was located in Victoria Land, (2) to determine the setting in which glaciogenic sediments were deposited (i.e., subglacial, periglacial, glaciomarine), and (3) to document the occurrence of glaciotectonic structures within upper Paleozoic strata in Antarctica. Understanding upper Paleozoic glaciogenic strata in southern Victoria Land will help to resolve the nature, timing, and extent of the LPGIA, which is of importance, as this glacial interval is one of the best analogues for predicting what will happen to Earth systems during the transition out of the current Cenozoic Ice Age. STRATIGRAPHY Glaciotectonic structures and glaciogenic deposits described in this paper are found at the top of the Devonian Aztec Siltstone, in the Lower Permian Metschel Tillite, and at the base of the Lower Permian Weller Coal Measures in southern Victoria Land (Fig. 3). The rocks are part of the Taylor (Devonian) and Victoria (Permian and Triassic) Groups of the Beacon Supergroup. Barrett (1972), Barrett and Kyle (1975), McKelvey et al. (1977), Barrett and McKelvey (1981), and McElroy and Rose (1987) provide

Figure 3. Generalized stratigraphic section of Devonian and Permian rocks exposed in southern Victoria Land, Antarctica.

an overview of the occurrence and distribution of Devonian and upper Paleozoic rocks in this area. The Aztec Siltstone is up to 217 m thick and consists of interbedded shale, siltstone, and cross-bedded sandstone (Fig. 3). Siltstones and shales are a few centimeters to several meters thick, whereas sandstones vary from 0.1 to 15 m in thickness (McKelvey et al., 1977; McElroy and Rose, 1987). The occurrence of fining-upward cycles, red beds, rootlet horizons, mud cracks, and conchostracan fossils all suggest that deposition occurred in an alluvial setting (McKelvey et al., 1977; McPherson, 1978, 1979). A microflora containing Geminospora lemurata and fossil fish near the top of the unit, including Bothriolepis, Groenlandaspis, and Turinia gondwana, indicate a Middle to Late Devonian age (Helby and McElroy, 1969; McKelvey et al., 1972; Ritchie, 1975; Young, 1988, 1989, 1991; Turner and Young, 1992).

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Glaciogenic strata of the Metschel Tillite unconformably overlie rocks of the Aztec Siltstone (Fig. 3). The Metschel Tillite is 0–85 m thick and consists of diamictite, conglomerate, sandstone, and shale, which are locally intraformationally deformed (Barrett and Kyle, 1975; McKelvey et al., 1977; Barrett and McKelvey, 1981; McElroy and Rose, 1987). Although the age of the Metschel Tillite is unknown, it is here considered to be Early Permian on the basis of its position below rocks of the Weller Coal Measures, which contain Lower Permian palynomorphs, and also on the basis of correlation of the Metschel Tillite with glaciogenic rocks of the Darwin Tillite exposed 130 km to the south. The Darwin Tillite contains Asselian-Sakmarian palynomorphs, as do all other palynomorph-bearing upper Paleozoic glaciogenic rocks in Antarctica (cf. Barrett and Kyle, 1975; Kyle, 1977; Kyle and Schopf, 1982; Lindström, 1995; Askin, 1998). Strata of the Weller Coal Measures rest both disconformably and conformably on diamictites of the Metschel Tillite (Fig. 3; McKelvey et al., 1977). Conglomerate and cross-bedded, coarsegrained sandstone containing quartz pebbles occur at the base of the formation, whereas upward within the unit, fine to coarsegrained sandstones are interstratified with siltstones, shales, and coals. Strata near the base of the formation contain fossil leaf impressions of Gangamopteris, Glossopteris, and Cordaites (Kyle, 1976; Pyne, 1984; Cúneo et al., 1993). Kyle (1977) correlated the Weller Coal Measures with Lower Permian (Artinskian) rocks in Australia (Australian Stage 4 Palynomorph Zone of Evans, 1969) based on microfloras from the middle and upper parts of the formation. GLACIOGENIC DEPOSITS AND GLACIOTECTONIC STRUCTURES Two types of deformed and two types of undeformed facies associations within Devonian and upper Paleozoic strata are reported in this paper. The associations are (1) sheared and homogenized sandstone and diamictite, (2) large-scale thrust sheets, (3) diamictite, and (4) sheet sandstone facies associations.

gently dipping (12°–24°) sandstone, mudstone, and paleosols in the Aztec Siltstone that are internally folded (folding may be the result of numerous microscopic thrust faults) and cut by thrust faults with centimeter-scale displacements; Fig. 4A); (3) 2 m of pervasively thrust faulted Aztec Siltstone where individual sandstone beds are internally folded and thrust faulted (Fig. 4B); (4) a sharp contact that truncates underlying folded and faulted rocks of the Aztec Siltstone below from diamictites of the Metschel Tillite above; (5) 1.5 m of thrust faulted diamictite consisting of a chaotic mixture of sandstone pods and boudins, including material from the Aztec Siltstone, and granite and quartz clasts up to 0.3 m in diameter; (6) 2–3 m of pervasively foliated (dips of 8°–30°), chaotic to homogenized diamictite containing granite and quartz clasts up to 0.2 m in diameter (Fig. 4C); and (7) locally up to 2 m of almost completely homogenized foliated diamictite containing granite, and quartz clasts up to 0.2 m in diameter (Fig. 4D). Foliation, thrust faults, and axial planes of folds within the deformed zone dip at up to 44° toward 306°. At the top of the deformed zone a sharp contact separates highly foliated and homogenized diamictites below from undeformed, weakly stratified to stratified diamictite above (Fig. 4E). A similar but much thinner deformed zone occurs on the east side of Mount Ritchie near the top of the Metschel Tillite. There, weakly stratified and stratified diamictite are overlain by a 1-m-thick sandy interval containing numerous small-scale thrust faults, and an overlying 1-m-thick foliated diamictite. The small-scale thrust faults dip at up to 45° toward 243°. This foliated diamictite is overlain by a sharp contact with weakly stratified diamictite or by an erosional surface that separates strata of the Metschel Tillite below from conglomerates and cross-bedded sandstones of the Weller Coal Measures above. Interpretation The presence of small-scale thrust faults, chaotic bedding– boudinage structures, and foliated homogenized materials at Mount Metschel suggest that sandstone and shale at the top of the Aztec Siltstone and diamictite at the base of the Metschel Tillite were deformed by shear. In general, deformation within

Sheared and Homogenized Sandstone and Diamictite Facies Association Description On Mount Metschel (Fig. 2) an asymmetric deformation zone separates the 50(+)-m-thick Devonian Aztec Siltstone from the overlying 21.5-m-thick Permian Metschel Tillite (Fig. 4). At this site the Aztec Siltstone consists of 0.1–0.5-m-thick interbeds of cross-laminated and horizontally laminated, fine- to mediumgrained sandstone, gray to red mudstone, and calcrete-bearing paleosols. Near the top of the unit, strata pass upward through a deformed zone, which shows increasing strain upward, into foliated diamictite of the Metschel Tillite (Fig. 4). The following zones occur from the base to the top of this sequence, which straddles the Aztec-Metschel contact: (1) undeformed interbeds of sandstone and mudstone of the Aztec Siltstone; (2) 2–2.5 m of

Figure 4. Stratigraphic column of deformed strata of the Devonian Aztec Siltstone and Permian Metschel Tillite exposed on the eastfacing slopes of Mount Metschel, interpreted to be the deposits of a subglacial deformation bed (glaciotectonite and deformation till). Orientations of structural features and interpreted ice-flow directions are also shown. (A) Thrust faults with centimeter-scale displacements cut strata of the Aztec Siltstone. Brunton compass for scale. (B) Pervasively thrust faulted sandstone at the top of the Aztec Siltstone. Rock hammer for scale. (C) Pervasively foliated diamictite and sandstone near the base of the Metschel Tillite. Ice axe for scale. (D) Pervasively foliated and partially homogenized diamictite near the base of the Metschel Tillite. Ice axe for scale. (E) Stratified diamictite with lobeshaped, boulder-bearing sandy diamictite near the top of the Metschel Tillite at Mount Metschel. Ice axe for scale. vm—vector mean, mvl— mean vector length, n—number of measurements.

E

D

Stratified diamictite

Dipping foliation

Lobe-shaped sandy diamictite

(c)

Ice flow direction

n=11

Sand f mc

vm=126 mvl=0.95 n=11

Diamictite

Clay Silt

Dip of shear planes

C

14m

Metschel Tillite

Undeformed stratified diamictite with granite and quartz clasts up to 1 m in diameter

(d)

B

Homogenized diamictite with granite and quartz clasts up to 0.2 m in diameter Pervasively foliated, diamictite with granite and quartz clasts up to 0.2 m in diameter Thrust faults

Thrust faulted diamictite with sand boudins and granite and quartz clasts up to 0.3 m in diameter

Sharp Contact Aztec Siltstone

Pervasively thrust faulted (dm to m displacements) sandstone that is internally folded and thrust faulted

A Small-scale thrust faults

Internally folded and thrust faulted with cm displacements Small-scale thrust faults

0

Undeformed Aztec Siltstone

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individual tectonic and subglacial shear zones displays a symmetrical increase in strain away from upper and lower boundaries toward a homogenized central region of high finite shear strain (Van der Wateren, 1987, 2002). The distribution of strain in the shear zone at Mount Metschel is asymmetrical, with changes in structures indicating that strain increased upward from undeformed rock in the Aztec Siltstone into the overlying homogenized diamictite at the base of the Metschel Tillite. However, structures indicating decreasing strain do not occur as the homogenized diamictite passes abruptly upward into undeformed stratified diamictite. Rocks at the contact between the Aztec Siltstone and the Metschel Tillite are similar to sediment deformed and deposited subglacially. Under certain conditions, shear between glacial ice and its unconsolidated or weakly consolidated substrate results in formation of a deforming bed beneath the ice-sediment interface (Alley, 1991; Benn and Evans, 1998). For ice streams and surging glaciers, much of the forward motion of the glacier may be accounted for within such a unit. Formation of a deforming bed is favored by the occurrence of (1) subglacial waters, (2) unfrozen unconsolidated or weakly consolidated (sedimentary) substrates, and (3) substrate materials that inhibit water from draining away into the underlying sediments (Alley, 1991; Boulton, 1996). Under these conditions, glacial ice forms the “mirrored top” of the idealized symmetrical shear zone with strain increasing upward within the substrate toward the ice-sediment interface and then decreasing away from the interface into the overlying glacier (Boulton, 1979; Boulton and Jones, 1979; Boulton and Hindmarsh, 1987; Van der Wateren, 2002). A vertical strain profile within a subglacial deforming bed consists of (1) undeformed sediment at depth, (2) materials deformed by brittle failure and faulting, (3) pervasive ductile shearing of materials, and (4) plowing and sliding of ice and debris along the ice-sediment contact (Banham, 1977; Boulton, 1987; Alley, 1991; Hart and Boulton, 1991; Benn, 1995; Benn and Evans, 1996; Van der Wateren, 2002). A décollement may separate structural zones displaying different degrees of strain (Banham, 1977; Boulton, 1987). Subglacially sheared deposits are classified by Benn and Evans (1996) as glaciotectonites if the deposits contain structural characteristics of the original parent material, or as deformation tills if the material has been homogenized. Recently, van der Meer et al. (2003) reported that all subglacial tills form from a combination of both deformation and lodgment. At Mount Metschel, deformed Aztec Siltstone has all of the characteristics of a glaciotectonite, whereas the overlying foliated diamictites at the base of the Metschel Tillite display characteristics of a deformation till. The two formations are separated by a décollement, which would have served as the structural boundary between brittle and ductile deformation. The décollement also occurs at the position of the regional unconformity that now separates the two formations. The occurrence of this type of deformation indicates that rocks of the Upper Devonian Aztec Siltstone were weakly consolidated during glaciation. Such an interpretation is tenable, as Cambrian and Ordovician sandstones

of the upper Midwestern United States are weakly consolidated and were locally deformed by glaciogenic activity during the Pleistocene. The interstratification of sands and muds in the Aztec Siltstone likely inhibited subglacial drainage into underlying aquifers, possibly resulting in high pore-water pressures and the formation of a subglacial deforming bed during overriding of the area by a Permian glacier. The orientation of the thrust faults and shear planes at Mount Metschel indicate local ice movement toward 126°. Barrett and Kyle (1975) reported striations at Mount Metschel oriented toward 120°. However, these striations were likely slickensides contained within the shear zone. The deformed zone at the top of the Metschel Tillite on the east-facing slope of Mount Ritchie is also interpreted to have formed subglacially as a glaciotectonite and deformation till. The occurrence of weakly stratified and stratified diamictite (see section below on Diamictite Facies Association) directly below this unit, coupled with the orientation of the structural fabric within the deformed zone, suggests that grounded ice advanced northeastward (063°) into a glaciomarine setting. Large-Scale Thrust Sheet Facies Association Description of Strata at Mount Richie At Mount Ritchie (Fig. 2), thrust sheets in the Metschel Tillite, up to 15 m thick, occur at and just above the contact with undeformed strata of the Devonian Aztec Siltstone (Fig. 5). The sheets, which are imbricately stacked, make up a 600+-m-long duplex along a north-northwest– to south-southeast–trending ridge near the summit of the mountain. These strata are exposed on the west-facing slope of Mount Richie. Along much of the ridge the contact between the thrust sheets and underlying strata is covered by snow and scree. Where exposed, the substratum includes (1) diamictite, (2) deformed interstratified diamictite and sandstone, (3) breccia, and (4) strata of the Aztec Siltstone (Figs. 5A, 5B, 5C). At its northern end the duplex ramps up onto a 4–6-m-thick, fine- to medium-grained massive sandstone (see the section below on Sandstone Sheet Facies Association) in the Metschel Tillite. There, individual thrust sheets are bounded by

Figure 5. (A) Photo of imbricate thrust sheets exposed on the westsouthwest face of Mount Ritchie, interpreted to be part of a subaquatic thrust moraine complex. For scale, the distance from the base of the Weller Coal Measures to the base of the sill is 25 m. (B) Interpretive map of the thrust sheets on the west-southwest face of Mount Ritchie. Stereonet shows the orientation of the thrust sheets, and the rose diagram shows paleo ice-flow directions interpreted from the orientations of the thrust sheets. (C) Brecciated zone marking the location of a shear plane at the base of a thrust sheet on Mount Ritchie; 15-cm-long ruler for scale. (D) Thrust sheets of the Metschel Tillite, composed of strata of the Aztec Siltstone at Mount Ritchie. The sheets are bounded by major thrust faults (MTF), whereas recumbent folds and small-scale thrust faults (STF) with meter-scale displacements occur within the sheets. Person and Jacob’s staff (1.6 m) for scale. vm—vector mean, mvl—mean vector length, n—number of measurements.

A

B

C

D

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listric to sigmoidal-shaped shear zones that consist of either a 0.05–0.3-m-thick, intensely foliated mudstone containing rare granite and gneiss granules, pebbles, and cobbles or locally by a 0.5–3-m-thick brecciated and/or chaotically folded zone. The thrust sheets are sigmoidal in shape and consist of massive sandstone, which contains an internal mosaic of annealed fractures and dewatering structures. The sheets within this complex dip at up to 45° toward 127°. Near the back of the thrust complex, several thrust sheets consist of an interbedded succession of crossstratified and horizontally laminated sandstones and mudstones of the Aztec Siltstone (Figs. 5A, 5B, 5D). These sheets rest on the backs of the thrust sheets described above. Internally, the Aztec thrust sheets contain internal deformational structures that include (1) high-angle normal faults, (2) small-scale reverse faults with displacements of centimeters to a few meters, (3) overturned folds, and (4) recumbent folds (Fig. 5D). The folds and smallscale thrust faults indicate compression toward the WNW. The thrust sheets are overlain by weakly stratified and stratified diamictite (see Diamictite Facies Association, below) that lap onto and drape over the thrust complex. Description of Strata at Kennar Valley At Kennar Valley (Fig. 2), rocks of the Metschel Tillite are exposed primarily on a ridge that extends northeastward into the center of the valley. Here, highly deformed rocks of the Metschel Tillite occur between undeformed strata of the Devonian Aztec Siltstone below, and undeformed strata of the Weller Coal Measures above (Figs. 6 and 7). Within the Metschel Tillite the style of deformation changes progressively along the ridge toward the southwest from ductile, to brittle, to undeformed (Fig. 6). A chaotic zone of ductile deformation is exposed on the eastern valley wall and on the northeastern end of the central ridge. The folding consists of large-scale recumbent and isoclinal folds and/or nappe-like structures, with structural displacement toward the west-southwest. The strata consist of interstratified, mediumto coarse-grained, massive to internally deformed sandstones and stratified diamictites (Figs. 6A–6D). Some of the sandstone units contain dewatering features (pipes, sheets, and dish structures). Low-angle thrust-fault–bounded packages up to 10 m thick occur throughout the middle and southern parts of the central ridge (Figs. 6 and 7A–7C). These thrust sheets occur in front, and on top of, the chaotic ductile zone described above. The faults, marked by 0.01–1-m-thick, boudin-bearing, foliated and homogenized mudstones and siltstones (Figs. 7D and 7E), dip at up to 51° toward the east-northeast (toward 72°), indicating structural transport of the sheets toward the west-southwest (252°). These faults have high-angle dips along the northeast part of the ridge but become subhorizontal toward the southwest. Internally within the thrust-fault–bounded packages, beds are slightly folded and contain abundant high-angle normal and listric-shaped reverse faults (Figs. 7C, 7F, 7G). The normal faults are common at the bases of the thrust sheets and occur within the concave-up troughs of small, open synclines (Fig. 7C). On the west-northwest side of the central ridge (Fig. 7A and 7B), stacked thrust sheets

1–7 m thick consist of (1) weakly stratified diamictite; (2) interbedded conglomerate and faulted, cross-bedded to massive sandstone containing dish structures and dewatering pipes; and (3) coarsening-upward siltstone to sandstone successions with siltstones containing isolated ripples (lenticular bedding) and sandstones containing cross-laminae, horizontal laminations, wave ripple laminae, swaley cross-stratification, load and flame structures, and scattered dropstones. The contact with the underlying Aztec Siltstone is covered by scree. However, some thrust sheets cut below the level of the contact (Figs. 7A and 7B). Undeformed conglomerates and sandstones of the Weller Coal Measures rest on an erosion surface with a relief of several meters cut into the underlying Metschel Tillite. Interpretation Deformational structures in the Metschel Tillite at Mount Ritchie and Kennar Valley indicate formation caused by horizontal compression. These structures include listric-shaped thrust faults, sigmoidal-shaped thrust sheets, fold and thrust nappes, and recumbent folds. Shear zones at the base of individual thrust sheets are marked by boudin-bearing, foliated mudstones and decimeter- to meter-thick breccia zones. Substantial evidence indicates that deformation occurred primarily during deposition of the Metschel Tillite. This evidence includes (1) occurrence of the thrust sheets between undeformed strata of the Aztec Siltstone and glaciogenic deposits of the Metschel Tillite below, and undeformed glaciogenic and fluvial strata of the Metschel Tillite and Weller Coal Measures above; (2) truncation and overriding of glaciogenic deposits of the Metschel Tillite by the thrust sheets; (3) composition of the thrust sheets, which consist primarily of glaciogenic deposits; (4) occurrence of both soft sediment deformational (i.e., massive sandstones, dewatering structures, and chaotically folded sandstones and diamictites) and brittle (normal and reverse faults) structures within the thrust sheets, indicating that both ductile and brittle deformation and water expulsion from the sediments occurred, which is characteristic of deformation of unconsolidated and weakly consolidated deposits; and (5) onlapping and overlapping of the thrust complexes by diamictites, which indicate that continued glaciogenic deposition occurred across relief generated by the compressional structures. Thrust sheets truncating (Kennar Valley), and composed of strata (Mount Ritchie) derived from the underlying Aztec Siltstone, indicate that some excavation of bedrock also occurred. Because the zone of deformation appears to have occurred in unconsolidated glaciogenic sediment, and because it is confined between undeformed strata of the Aztec Siltstone below and the Permian Weller Coal Measures above, the compressional structures are interpreted to be intraformational features associated with glacial activity rather than being due to tectonic stresses. In glaciogenic settings, compressional features result from either sliding-slumping or glaciotectonic deformation, both of which produce similar structures. Slides occur on slopes where gravitational failure of the substrate results in rotational

A

Weller Coal Measures

B Brittle zone Unconformities Surfaces (bedding and faults)

Chaotic zone (Fig. 6C)

Metschel Tillite

C

D

E Brittle

Ductile

Glacier

Figure 6. Folded and thrust-faulted strata of the Metschel Tillite at Kennar Valley that are interpreted to have formed as a subaquatic thrust moraine. (A, B) Highly folded strata (right side of photo), giving way to strata deformed by thrust faults (left side of photo). The strata are exposed on the east-northeast end of the central ridge in Kennar Valley. The photo shows a slope and cliff face ~50 m high. (C) Closeup of the highly folded portion of the cliff face shown in Figures 6A and 6B. (D) Highly deformed strata contained within the chaotic zone exposed on the westnorthwest side of the central ridge in Kennar Valley. Jacob’s staff (1.6 m) for scale. (E) Schematic diagram showing formation and orientation of deformation in a modern thrust moraine (diagram modified from Croot, 1988).

A

B

C

D Thrust sheet

Major thrust fault Normal faults

Thrust fault

E

F

Major thrust fault

G

Major thrust fault

Normal faults

Internal thrust fault

Figure 7. Strata exposed on the west-northwest–facing slope of the central ridge in Kennar Valley, interpreted to be part of a thrust moraine complex. (A, B) Thrust faults and thrust sheets exposed along the northern end of the central ridge. The Weller Coal Measures are ~28 m thick, for scale. Stereonet shows the orientation of thrust faults, and the rose diagram shows the ice-flow directions interpreted from fault orientations. (C) Thrust sheet showing basal thrust fault and internal accommodation faults. Jacob’s staff (1.6 m) for scale (white arrow). (D, E) Foliated and homogenized mudstones mark the position of major thrust faults. (F) Small-scale thrust fault at the base of a thrust sheet. Jacob’s staff (1.6 m) for scale (white arrow). (G) Normal faults near the base of a thrust sheet; 15-cm-long ruler for scale. VM—vector mean, MVL—mean vector length, N—number of measurements.

Permian deposits, southern Victoria Land, Antarctica extension of a sediment mass along listric glide planes followed by downslope sliding of a coherent block away from a headwall. Deposition occurs when the block comes to rest at the toe of the slide. Slumps form in a similar fashion. However, slumps display internal folding and faulting. The resulting slide-slump structures consist of either a single block or stacked blocks or sheets (multiple slide-slump events), each bounded below by a sheared glide plane (Allen, 1985; Collinson and Thompson, 1989; Ricci Lucchi, 1995; Benn and Evans, 1998). Slump and/or slide blocks are identified by the following criteria: (1) truncation of underlying strata from rotational extension in areas adjacent to the slumpslide scarp; (2) deposits, which at the head of the slump-slide dip away from the headwall scarp; (3) deposits, which at the toe of the slump-slide typically rest on slopes that dip in the direction of sliding; (4) deposits typically composed of only one to several sheets with younger sheets resting on the backs of older sheets; (5) the occurrence of slump fold noses; (6) an increase in the degree of deformation in the direction of transport owing to compression at the toe of the slump-slide; and (7) deposits that typically do not contain excavated bedrock blocks (cf. Tucker, 1996; Collinson and Thompson, 1989; Ricci Lucchi, 1995; Stow, 2005). Slides and slumps are commonly associated with debris flows, which are often triggered during formation and movement of the slump and slide blocks. Structures formed by glaciotectonic compression develop at ice margins during glacial advance and involve displacement of subglacial and proglacial sediment, and weak rock by both ductile and brittle deformation (Aber et al., 1989; Hart and Boulton, 1991). Deformation is facilitated by (1) horizontal stresses resulting from high ice overburden pressures up-glacier, (2) weak subglacial and periglacial substrates that produce décollements during failure, (3) high pore-water pressures that reduce the cohesive and yield strength of ice marginal and subglacial sediments and coherent substratum, and (4) by the presence of reverse slopes along the glacial margin that increase horizontal stress (Bluemle and Clayton, 1984; Aber et al., 1989; Boulton and Caban, 1995). Ductile deformation includes formation of open to overturned folds, whereas brittle failure results in the development of lowangle overthrusts, listric thrust faults, imbricately stacked thrust sheets, and nappes (Hart, 1990; Van der Wateren, 2002). “Piggyback” thrusting is common where early formed proximal sheets are carried on the backs of later formed distal blocks. In this scenario, intense folding and high-angle reverse faulting occur along the ice margin, decreasing to lower angle emplacement of thrust sheets outward (Eybergen, 1987; Croot, 1988; Van der Wateren, 2002). Thrust moraines are the surface expression of this deformation (Aber et al., 1989). Several criteria can be used to identify thrust sheet complexes produced by glaciotectonic deformation. These include (1) truncation of underlying strata throughout the zone of deformation, especially along frontal thrusts; (2) occurrence of compressional features throughout; (3) décollement–thrust fault surfaces that ramp up and over truncated proglacial sediment near the leading edge of the frontal thrust; (4) thrust complexes where older

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thrust sheets rest on top of younger sheets or where early formed sheets display near-vertical orientations owing to rotation and piggyback transport on younger sheets that formed along frontal thrust faults during continued compression; (5) a decrease in the degree of deformation in the direction of transport owing to the formation of frontal thrust faults; and (6) thrust complexes that typically contain excavated blocks of bedrock (Croot, 1988; Aber et al., 1989). Small thrust sheets composed of periglacial material also occur along both terrestrial and subaqueous ice margins. These thrust blocks or push moraines, which are typically no more than 5 m high, form during small-scale seasonal advance and retreat cycles of the ice front. During a seasonal advance, periglacial material is pushed or “bulldozed” into a morainal ridge. Although these features are commonly composed of supraglacial debris dumped at the ice margin, some also include proglacial sediment (Bennett and Glasser, 1996). Compressional deformation in the Metschel Tillite is consistent with formation by glaciotectonic deformation of periglacial deposits. Supraglacial material does not occur within the thrust sheets. At Mount Ritchie, shear zones truncate glaciogenic and/or Devonian strata throughout the thrust complex. Near the front of the complex the basal thrust sheet ramps up and over preexisting deposits, with the orientation of the truncation surface dipping in a direction opposite to that of the direction of transport for the sheets. Because thrust sheets of the Aztec Siltstone rest on these basal imbricated sheets, the “Aztec” sheets likely formed early, only to be later transported piggyback on younger structures that developed along frontal thrust faults. Dip directions of the thrust sheets at Mount Ritchie suggest glaciotectonic transport, and therefore glacial advance toward 307°. At Kennar Valley the lateral change from highly contorted beds, recumbent folds, and thrust nappes to thrust sheets indicates a decrease in the intensity of deformation from ductile to brittle in the direction of structural transport, which was toward 252°. The high-angle orientation of the proximal deformation zone suggests that these may have been the first sheets to have formed and that they were later transported piggyback on younger thrust sheets and rotated into high angles during continued propagation of the thrust front. This progression of structures is similar to proximal to distal changes in deformation style seen within modern ice-marginal thrust moraine complexes (Fig. 6E; cf. Croot, 1988; Van der Wateren, 2002). Therefore, structures exposed at Kennar Valley are interpreted to have formed as glaciotectonic features along an ice margin. During emplacement, flexure of the deforming mass would have produced both extensional and compressional stresses within individual thrust sheets where the thrust sheets were flexed during transport into broad synclines and anticlines. Such stresses would have resulted in the formation of internal accommodation features (i.e., normal and listric-shaped reverse faults), which would have facilitated flexing of the thrust sheets. In modern glacial settings, glaciotectonic deformation occurs along both terrestrial and glaciomarine ice margins (Bennett

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et al., 1999; Van der Wateren, 2002). Massive, weakly stratified, and stratified diamictite facies associated with the thrust sheets at Mount Ritchie suggest that deposition occurred in a basinal setting (see next section). The lithologies and sedimentary structures of strata contained within the thrust sheets at Kennar Valley, which include stratified to massive diamictite (see next section), coarsening upward siltstone to sandstone successions containing wave ripple laminae, and swaley cross-stratification, also suggest that ice advanced into a glaciomarine environment. Diamictite Facies Association Description Massive to weakly stratified diamictite occurs at Kennar Valley and at Mount Ritchie (Figs. 2 and 8). At Kennar Valley several of the thrust sheets are composed entirely of this type of diamictite (see large-scale thrust-sheet-facies association, Figs. 8A and 8B). At Mount Ritchie, massive and weakly stratified diamictites typically have gradational upper and lower contacts with stratified diamictite. However, weakly stratified diamictite is also truncated by conglomerates and sandstones in the overlying Weller Coal Measures. The diamictite consists of a clay to fine-grained sand matrix containing scattered to abundant clasts. Laminae and thin wisps of silt and sand, which are better sorted than the surrounding matrix, define faint stratification. Clasts of granite, quartzite, and gneiss up to 1 m in diameter, pierce stratification, and the long axes of some clasts are oriented at high angles to bedding. Within massive to weakly stratified diamictite units, isolated centimeter- to meter-thick and decimeter- to meter-wide pods of contorted and massive sandstone and conglomerate with dewatering structures occur (Fig. 8C and 8D). Injection structures in the form of diamictite diapirs intrude these pods. Stratified diamictites are common at Mount Metschel, Mount Ritchie, and within the thrust sheets at Kennar Valley (Figs. 8E–8G). These diamictites have a similar matrix and clast (lithology and size) composition as those of the massive and weakly stratified diamictites. Stratification is distinct, however, and consists of centimeter- to decimeter-thick semi-continuous layers of silty sand–rich layers alternating with mud-rich layers. In a few places, intercalations of centimeter-scale units of normally graded to cross-laminated, medium-grained sandstones delineate stratification (Fig. 8G). The bases of the sandstones commonly display load structures. Within the stratified diamictite, clasts up to 30 cm in diameter are common and cut stratification. Within the stratified diamictites, decimeter-thick and tens-of-meters-long lenses, or lobe-shaped bodies of massive sandstone and conglomerate, occur (Fig. 4E). These bodies characteristically contain sharp bases with abundant centimeter- to decimeter-scale load structures. The bodies are also intruded by diapirs of stratified diamictite. Boulders up to 0.7 m in diameter protrude from the top of the lobe-shaped bodies. On the north-northwest and west-southwest sides of Mount Ritchie, just below the contact with the overlying Weller Coal Measures, locally abundant accumulations of clasts up to 1 m

in diameter occur (Fig. 9A). Many of these clasts cut stratification and have their long axes oriented at high angles to bedding. Faceted and striated clasts also occur. At the base of the overlying Weller Coal Measures, gravel clasts (pebbles, cobbles, and boulders) protrude downward into the underlying diamictites, as do load structures at the base of coarse-grained sandstones and pebble to cobble conglomerate units. Intrusion of the Weller by diamictite diapirs also occurs, as does interfingering of Metschel diamictites and Weller sandstones and conglomerates (Fig. 9B). These coarse-grained Weller sediments are part of a 6- to 10-m-thick wedge-shaped coarse- to very coarse grained trough-cross-bedded sandstone body (Figs. 10A–10D). When traced over several hundred meters, this body displays an internal geometry characterized by low-angle downlapping beds and surfaces. When viewed on exposures perpendicular to paleoflow orientations, the downlapping units are cut by numerous smallscale (meters to tens of meters wide and up to a few meters deep) cross-bedded, sand-filled channels. Paleocurrent directions are oriented toward 234° (Figs. 10C and 10D). Interpretation Massive and stratified diamictites are commonly interpreted to have formed in both polar and temperate glaciomarine systems by a number of different processes. In polar settings the temperature and strength of the ice allow for the development of floating glacial tongues and ice shelves. Under these conditions, ice decouples from the bed and begins to float just seaward of a grounding line. Here, in this proximal setting, melt-out of basal debris from the underside of the glacier allows both fine- and coarse-grained particles to settle through the water column and to produce massive diamictites. Sedimentation rates are relatively high near the grounding line, but they decrease distally owing to the loss of debris-rich basal ice (Evans and Pudsey, 2002). Stratified diamictites occur where sedimentation rates are lower, and where bottom currents (e.g., tidal pumping) winnow out finer grained particles (Domack et al., 1999). Dropstones are produced by either iceberg rafting in open-marine settings or as rain-out from the debris-poor distal portions of the ice shelf or glacial tongue (Evans and Pudsey, 2002). Under temperate conditions, warmer, weaker ice typically results in tidewater glaciers entering the sea as an actively calving ice cliff. Under these conditions the glacial terminus is grounded or partially floating. Here, glacial deposition is dominated by meltwater and, to an equal or lesser extent, by rafting of debris by icebergs. Grounding-line fans form where subglacial meltwaters, emerging from conduits along the ice front, deposit sand and gravel. As velocity in the effluent flow drops, buoyant forces become dominant, and the inflowing fresh water rises to the surface to form an overflow plume that transports fine sand, silt, and clay seaward. Sedimentation occurs as particles released from the plume settle through the water column (Cowan and Powell, 1990). Plume sedimentation, in conjunction with dumping of coarse debris during calving at the ice front, and by the release of clasts from the melting of icebergs, produces stratified

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Figure 8. (A) Massive diamictite contained within two different thrust sheets exposed on the west-northwest–facing slope of the central ridge in Kennar Valley. Person for scale. (B) Massive diamictite grading into weakly stratified diamictite on the west-northwest–facing slope of the central ridge in Kennar Valley. The diamictite unit is ~4 m thick at its thickest point in the photo. (C) Deformed sandstone lens contained within weakly stratified diamictite exposed on the eastern side of Mount Ritchie. Person for scale. (D) Deformed pebbly sandstone lens containing diamictite diapirs near the summit of Mount Ritchie. Trekking pole for scale. (E) Stratified diamictite containing a lens of massive diamictite at Mount Metschel. Ice axe for scale. (F) Stratified diamictite, Mount Ritchie; 15-cm-long ruler for scale. (G) Interstratified diamictite and beds of medium- to coarse-grained sandstone, Mount Ritchie. Persons for scale.

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Figure 9. (A)Weakly stratified diamictite at the top of the Metschel Tillite at Mount Ritchie. Note the abundance of clasts and the high dip angle of the long axes of many of the clasts. Clasts in the overlying Weller Coal Measures also project down into the underlying diamictite. Hammer for scale. (B) Interfingering of weakly stratified diamictite of the Metschel Tillite, and sandstone and conglomerate of the Weller Coal Measures. Trekking pole for scale.

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Figure 10. (A, B) Sandstone sheet containing diamictite lenses and a channel body in the Metschel Tillite. A wedge-shaped sandstone at the base of the Weller Coal Measures containing downlapping surfaces is also shown. For scale, the distance from the base of the Weller Coal Measures to the base of the dolerite sill (Jurassic) is 25 m. (C, D) Wedge-shaped sandstone body at the base of the Weller Coal Measures, containing broad, channel-like scours. For scale, the distance from the base of the Weller Coal Measures to the base of the dolerite sill is 25 m. (E) Dewatering pipes within massive sandstone of the sandstone sheet facies association. Mechanical pencil for scale.

Permian deposits, southern Victoria Land, Antarctica diamictites in proximal glaciomarine positions (Cowan and Powell, 1991; Smith and Andrews, 2000). Massive diamictites occur in ice distal settings owing to iceberg rafting and scouring (Dowdeswell et al., 1994). However, massive and stratified deposits and/or mudstones lacking clasts can occur in either proximal or distal glaciomarine positions because of changes in sea ice and/or oceanographic conditions (Smith and Andrews, 2000; Dowdeswell et al., 2000). In both polar and temperate settings, stratification is also produced by periodic remobilization of the deposits by sediment gravity flows (Evans and Pudsey, 2002; Powell and Domack, 2002). In the Metschel Tillite, the occurrence of massive, weakly stratified, and stratified diamictites with gradational upper and lower contacts is suggestive of deposition in proximal to distal glacial basinal settings during fluctuations in the location of the ice front. Although it is unknown whether strata in southern Victoria Land were deposited under glaciolacustrine or glaciomarine conditions, a glaciomarine setting seems likely owing to an abundance of stratified diamictites, which suggests that sedimentation occurred from settling of particles from buoyant low-density meltwater plumes. Clasts in these deposits were likely introduced from the melting out of debris from ice fronts or from the release of debris rafted by icebergs. The abundance of massive, normally graded, and cross-bedded sandstone layers and pods within the strata is indicative of deposition from subaqueous meltwater as tractive flows and as sediment gravity flows, suggesting temperate glacial thermal conditions (Mackiewicz et al., 1984; Powell and Domack, 2002). Sandstones were deposited subaqueously on water-saturated substrates, which, when loaded, failed, producing load structures, dewatering structures, and disruption of the sandstone bodies by intrusion of diamictite diapirs. Such unstable substrates are the result of high sedimentation rates in ice proximal zones (Boulton, 1990). The occurrence of lens- and lobe-shaped sandy diamictite beds with protruding boulders is suggestive of deposition from debris flows with the boulders introduced as ice-rafted debris. However, small lenses of sand and conglomerate could have formed as iceberg dump structures (Thomas and Connell, 1985). Ice rafting of debris was likely an important component of sedimentation in both proximal and distal locations as indicated by the occurrence of clasts penetrating stratification (Thomas and Connell, 1985). Locally abundant boulders at the top of the Metschel Tillite at Mount Ritchie may indicate either iceberg dump structures or dumping of clasts at the ice front during calving events (Thomas and Connell, 1985; Powell and Domack, 2002). Further evidence of a glacial origin for these strata is provided by the occurrence of striated and faceted dropstones. Strata at the base of the overlying Weller Coal Measures show evidence of deposition contemporaneous with or shortly following deposition of the Metschel Tillite. Evidence includes (1) interfingering of Metschel diamictites with Weller sandstones and conglomerates, (2) boulders and cobbles in the Weller conglomerates protruding downward into the underlying diamictite,

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and (3) Metschel diamictites intruding Weller strata as diapirs. The occurrence of load structures and diapirs suggests failure of a water-saturated Metschel substrate owing to deposition of the overlying Weller sandstones and conglomerates. Downlapping surfaces within the basal Weller strata indicate progradation of a wedge-shaped body across proximal Metschel glaciomarine deposits (massive diamictite and abundant dropstones). The basal sandstone and conglomerate body has a geometry and internal features that are similar to grounding-line fans described by Powell (1990) and Powell and Alley (1997). Sandstone Sheet Facies Association Description On the northwest side of Mount Ritchie, strata near the middle of the Metschel Tillite consist of a 16.5-m-thick succession of fine- to medium-grained cross-bedded sandstone, fine- to medium-grained massive sandstone containing dewatering pipes (Fig. 10E), shale, and massive to weakly stratified diamictite. These units are contained within a sheetlike sediment body that is in erosional contact with underlying shales and diamictites (Figs. 10A–10D). Sandstone within the sheet is laterally continuous. However, diamictite and shale are discontinuous within the sheet and either drape underlying beds or form lens- to podshaped bodies. On the eastern end of the west-northwest face of Mount Ritchie a large channel is incised to a depth of 10 m into the sheet (Figs. 10A and 10B). Laterally, channel margins are concordant with beds in the underlying sandstone. The channel is filled with clast-supported conglomerate, cross-bedded sandstone, massive sandstone containing dewatering pipes, and by massive to weakly stratified diamictite. The sheet and channel are overlain by, and interfinger with, massive, weakly stratified, and stratified diamictite (see section on Diamictite Facies Association, above). On the west-southwest side of Mount Ritchie the sandstone sheet is overridden by thrust sheets (see section on Large-Scale Thrust Sheet Facies Association, above; Figs. 10C and 10D). Interpretation The interfingering of sandstone and massive to stratified diamictite (see section on Diamictite Facies Association, above) suggest that the sandstone sheet was deposited in a glaciomarine setting. Features within the sandstone sheet, which include channeling, conglomerates, and cross-bedded sandstone, indicate deposition from high-energy tractive currents. Massive sandstone with dewatering pipes suggests rapid sedimentation rates possibly from either suspension or sediment gravity flows. These conditions are common near the grounding lines of temperate tidewater glaciers (Powell, 1990). In this setting, fresh-water effluent flow, emanating from the base of the glacier, deposits wedgeshaped bodies of sand and gravel known as grounding-line fans. These bodies are laterally continuous for hundreds of meters and are commonly cut by channels as the effluent flow cuts into early

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formed deposits (Powell, 1990). Away from the glacial front, current velocity drops, allowing the low-density fresh-water flow to detach from the bottom and rise to the surface to form a buoyant overflow plume. Rapid sedimentation from the rain-out of sand, silt, and clay from the plume results in deposits with high initial water contents (Cowan and Powell, 1990). These ice proximal deposits are unstable and are highly susceptible to dewatering and/or to remobilization as sediment gravity flows. Addition of coarse debris from ice rafting results in a complex interfingering of sands, conglomerates, muds, and diamictites (Powell, 1990; Powell and Domack, 2002). The sandstone sheet facies association is here interpreted as a grounding-line fan. This interpretation is consistent with the observed sedimentary features and with the overriding of the sandstone sheet on the west-southwest side of Mount Ritchie by ice proximal thrust sheets, which also are common near the grounding line of some tidewater glaciers (Bennett et al., 1999). OVERALL DEPOSITIONAL SETTING Deformational and depositional lithofacies associations in the Metschel Tillite suggest that sedimentation occurred in ice marginal and glaciomarine settings. Deformational features are interpreted as glaciotectonites, deformation tills, and thrust duplexes. Glaciotectonites and deformation tills resulted from subglacial deforming beds, which, in modern settings, typically form beneath ice streams, outlet glaciers, tidewater glaciers, and surging glacial lobes where high pore-water pressures facilitate deformation of unconsolidated or weakly consolidated substrates. Thrust duplexes formed as periglacial thrust moraines along ice margins. The occurrence of these strata and their glaciotectonic structures suggests that deposition in southern Victoria Land occurred at or near glacial termini. The occurrence of massive and stratified diamictites, lonestone-bearing deposits, sheet sandstones, and sandstones with swaley cross-stratification and wave ripple laminations suggests that Permian glaciers in southern Victoria Land advanced into, and retreated from glaciomarine settings. Owing to the occurrence of meltwater plume deposits (stratified diamictites), the depositional setting was most likely glaciomarine. Evidence in the form of glaciotectonite, deformation till, and cross-bedded sandstone indicates that abundant meltwater was present at the time of deposition, and therefore that temperate thermal conditions characterized the depositing ice. Such thermal conditions are characteristic of tidewater glaciers where deposition is dominated by subglacial deforming bed conditions, grounding line processes, meltwater outflow, meltwater plumes, and iceberg rafting of debris. Only a few ice-flow directions have been reported from upper Paleozoic glaciogenic rocks in southern Victoria Land (Barrett and Kohn, 1975; McKelvey et al., 1977). Flow directions from these data are ambiguous. Although deformation structures reported in this paper represent local glaciotectonic displacement, these features also provide a record of paleo-ice-flow directions. Transport directions were derived from displacement directions

of thrust sheets, orientation of foliation, and analyses of folds. Interpretation of these data suggests that glacial ice converged on southern Victoria Land off the East Antarctic craton to the west (e.g., glaciotectonite and deformation tills at Mount Metschel and Mount Ritchie) and off an area in the direction of the present Ross Sea to the east (Fig. 11; e.g., thrust sheets at Kennar Valley and Mount Ritchie). If these directions are correct, then ice advanced from glacial centers on opposite sides of the depositional basin. Expansion of ice from these centers then allowed advance of ice margins into a glaciomarine setting in southern Victoria Land. DISCUSSION Models for the Late Paleozoic Ice Age place Victoria Land beneath the center of an immense Gondwanan Ice Sheet that waxed and waned throughout the Mississippian, Pennsylvanian, and Permian (e.g., Scotese et al., 1999). These models predict that terrestrial ice flowed southward across southern Victoria Land out of a major glacial spreading center (Lindsay, 1970; Barrett, 1991; Veevers, 2001). However, lithofacies and paleocurrent data presented here do not support such a conclusion. Instead, the results of this study suggest that expansion of temperate glaciers, flowing out of smaller glacial centers, converged on southern Victoria Land and extended into and retreated out of a glaciomarine setting during the Early Permian. No evidence for Carboniferous glaciation occurs. Therefore, glaciogenic strata of the Metschel Tillite are inconsistent with deposition in this area from a single, massive, long-duration Gondwanan Ice Sheet. Recent work in other areas of Antarctica and Gondwana are also challenging the prevailing view of the extreme size and duration of late Paleozoic glaciation. In the central Transantarctic Mountains, ongoing facies and paleocurrent analyses do not support the traditional view of terrestrial ice flowing radially out of Victoria Land (Isbell et al., 1997, 2005, 2008; Isbell, 1999). Instead, results show that ice converged on an elongate basin whose long axis was oriented parallel to the present trend of the mountain range (Figs. 1 and 11). In the central Transantarctic Mountains, glaciers, grounded along basin margins in the direction of the present polar plateau and in the direction of the Ross Sea–Marie Byrd Land, flowed transversely off the margins into a glaciomarine setting. This scenario is similar to that proposed for southern Victoria Land, thus suggesting the occurrence of multiple Permian glacial centers in Antarctica (Isbell et al., 1997, 2005, 2008; Isbell, 1999). Elsewhere in Antarctica, interpretations of earlier work suggest that a glacial center on the Ellsworth Mountains crustal block supplied ice to glaciomarine settings in the Ellsworth and Pensacola Mountains (Fig. 11C; Frakes et al., 1971; Matsch and Ojakangas, 1991; Collinson et al., 1994). In Queensland and New South Wales, Australia, Jones and Fielding (2004), Birgenheier et al. (2005, 2009), and Fielding et al. (2005, 2008) reported the occurrence of short, discrete intervals of mountain-valley glaciations, ice caps, and/or small

Permian deposits, southern Victoria Land, Antarctica

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Figure 11. Tectonic transport direction for rocks of the Metschel Tillite as indicated by the orientation of (A) thrust faults and foliation in glaciotectonite and deformation till (Mount Metschel and Mount Ritchie) and the orientation of (B) thrust faults associated with thrust sheets (Kennar Valley, Mount Ritchie). (C) Map showing interpreted flow directions in southern Victoria Land, plotted with data from glaciogenic strata elsewhere in the Transantarctic and Ellsworth Mountains (EM), and in South Africa. Data from Frakes et al. (1971), Barrett (1981), Collinson et al. (1994), Visser (1997), Isbell (1999), Lenaker (2002). v.m.—vector mean.

ice sheets. Their findings are in marked contrast with earlier reports that suggested that much of Australia was covered by a continental-scale polar ice sheet. The hypothesis that numerous ice centers occurred in Gondwana during the late Paleozoic is not a new concept. Work by Crowell and Frakes (1970), Caputo and Crowell (1985), Eyles (1993), López Gamundí (1997), Limarino et al. (2002), Isbell et al. (2003), Henry et al. (2008), and Isbell et al. (2008) showed that multiple ice sheets, ice caps, and alpine glaciers diachronously waxed and waned as Gondwana drifted across the late Paleozoic South Pole. The glacial record in southern Victoria Land is consistent with the concept of multiple ice centers within Gondwana, and interpretation of the record disproves that Antarctica was covered by a single, massive ice sheet during the late Paleozoic. Identification of the size and duration of Gondwana glaciation is of great importance in developing an understanding of

Earth systems during the late Paleozoic. Data obtained from the Metschel Tillite clearly show that multiple glaciers were active in southern Victoria Land during the Permian. These results, coupled with recently reported data in Australia and South America, strongly suggest that Gondwana glaciation was characterized by numerous glacial centers and alpine glaciers rather than by a single massive ice sheet. Because the geographic area of ice cover, ice volume, and the number of ice sheets are directly related, multiple glaciers, for a given land area, contain considerably less ice volume than a single massive ice sheet (Crowley and Baum, 1991; Isbell et al., 2003). Therefore, multiple Gondwana glaciers would have had a completely different impact on Earth’s natural systems than that of a massive ice sheet. For example, the waxing and waning of multiple ice sheets would have resulted in considerably smaller changes in eustatic sea level than those produced by growth and decay of a single glacier covering the same geographical area.

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CONCLUSIONS Glaciotectonic deformation and glaciomarine deposits within the Metschel Tillite indicate that late Paleozoic glaciation was characterized by temperate glacial conditions along Permian ice margins in southern Victoria Land rather than by polar glacial conditions associated with a glacial spreading center as previously hypothesized. Deposition occurred beneath glaciers undergoing deforming bed conditions and in a glaciomarine setting in front of either alpine glaciers, outlet glaciers, ice streams, or surging glacial lobes. The direction of glaciotectonic transport indicates that ice converged on southern Victoria Land from two directions. One source of ice was from the direction of the present East Antarctic craton. The other source of ice was from the opposite direction, that of the present Ross Sea. These findings imply that multiple glaciers were active in this region. Because multiple ice sheets covering a given area contain less ice than a single ice sheet, multiple glaciers would have influenced Earth systems (e.g., eustatic sea level, climate) differently than the massive-ice-sheet models predict. ACKNOWLEDGMENTS Discussions with Rosemary Askin, Jim Collinson, Ellen Cowan, Dyana Czeck, Pete Flaig, Tom Hooyer, Mark Johnson, Carlos Oscar Limarino, and Molly Miller are greatly appreciated. I also thank Paul Lenaker, Rosemary Askin, Tim Cully, Molly Miller, and Keri Wolfe for their help in the field. Peter Barrett, Luis Buatois, Jim Collinson, Chris Fielding, and Antonio Carlos Rocha-Campos provided valuable comments on earlier drafts of this paper. The U.S. National Science Foundation, Raytheon Polar Services, the New York Air National Guard, Kenn Borek Air LTD, Trans World Logistics, and Petroleum Helicopters Incorporated provided logistic support for fieldwork in Antarctica. National Science Foundation grants OPP-9909637, ANT-0440919, ANT0635537, and OISE-0825617 supported this work. REFERENCES CITED Aber, J.S., Croot, D.G., and Fenton, M.M., 1989, Glaciotectonic Landforms and Structures: Dordrecht, Kluwer Academic Publishers, 200 p. Allen, J.R.L., 1985, Principles of Physical Sedimentology: London, George Allen & Unwin, 272 p. Alley, R.B., 1991, Deforming-bed origin for southern Laurentide till sheets?: Journal of Glaciology, v. 37, p. 67–76. Askin, R.A., 1998, Floral trends in the Gondwana high latitudes: Palynological evidence from the Transantarctic Mountains: Journal of African Earth Sciences, v. 27, p. 12–13. Banham, P.H., 1977, Glacitectonics in till stratigraphy: Boreas, v. 6, p. 101– 105. Barrett, P.J., 1972, Late Paleozoic glacial valley at Alligator Peak, southern Victoria Land, Antarctica: New Zealand Journal of Geology and Geophysics, v. 15, p. 262–268. Barrett, P.J., 1981, History of the Ross Sea region during the deposition of the Beacon Supergroup 400–180 million years ago: Journal of the Royal Society of New Zealand, v. 11, p. 447–458. Barrett, P.J., 1991, The Devonian to Jurassic Beacon Supergroup of the Transantarctic Mountains and correlatives in other parts of Antarctica, in

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The Geological Society of America Special Paper 468 2010

Formation of euxinic lakes during the deglaciation phase in the Early Permian of East Africa Thomas Kreuser Gebretinsae Woldu Geology Department, University of Asmara, P.O. Box 1220, Eritrea

ABSTRACT The continental glaciation of Gondwanaland in the Late Carboniferous–Early Permian left traces in many places in southern and eastern Africa. This paper focuses on the last glacial advance and consecutive deglaciation leading to the formation of large euxinic lakes with high concentrations of organic matter. The Idusi Formation in the Tanzanian Ruhuhu Basin (initiating the Karoo cycle, which extends into the Triassic) provides the type section for this depositional sequence. It is subdivided into a lower Lisimba Member, the basal unit of glacial origin, and an upper Lilangu Member, characterized by postglacial black shale and rhythmites as evidence of a climatic amelioration on a large regional scale in Africa. Thickness and facies variations are attributed to a pronounced paleotopography as the result of scouring glaciers and local tectonic events. There is a gradual change between the members, reflecting a continuous climatic amelioration and change of sediment supply. The lacustrine environment was terminated by the onset of braided stream deposition (Mpera Sandstone Member); an erosional unconformity between the units marks the start of initial rifting in the Early Permian. This is followed by the development of extensive coal swamps in a temperate climate, where organic matter predominated over clastic supply. Periglacial deposits with tillites and rhythmites, containing dropstones, are overlain by glaciolacustrine laminites intercalated with glaciofluvial marginal deltaic sediments. Deglaciation provided water and accommodation space for the evolution of extensive anaerobic stratified lakes, which were the focus of prolific deposition of organic matter. This black shale may contain up to 11% TOC (total organic carbon) content. Eventually, the lake became shallower and was succeeded by alluvial fan deposition. The duration of the glaciation and deglaciation was ~20–25 m.y., and the lacustrine phase lasted ~4–5 m.y. These ages have been verified by palynology (Granulatisporites confluens Oppel zone). The hydrocarbon potential of the black shale was estimated by Rock-Eval pyrolyses. Hydrogen index, maximum temperature (Tmax), and vitrinite reflection were used to determine kerogen type, maturity stage, and subsidence history. A promising potential with respect to gaseous hydrocarbon generation was detected from both the euxinic black shale and the overlying coals. A comparison with other Tanzanian

Kreuser, T., and Woldu, G., 2010, Formation of euxinic lakes during the deglaciation phase in the Early Permian of East Africa, in López-Gamundí, O.R., and Buatois, L.A., eds., Late Paleozoic Glacial Events and Postglacial Transgressions in Gondwana: Geological Society of America Special Paper 468, p. 101–112, doi: 10.1130/2010.2468(04). For permission to copy, contact [email protected]. ©2010 The Geological Society of America. All rights reserved.

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Kreuser and Woldu Karoo basins reveals similar conditions in TOC values and temperature history. The wide regional extent of the anaerobic lacustrine black shale of the deglaciation event in several eastern and southern African basins evinces a similar climatic and regional tectonic framework in the pre-breakup phase of Gondwanaland during the Early Permian. This period of time may be of some importance in the future when the economic potential with respect to hydrocarbon generation of the Permian basins is scrutinized in more detail.

INTRODUCTION The present work is a general review of the results of a working group from the University of Cologne, Germany, and the University of Dar es Salaam, Tanzania, during a research period of more than 15 yr. Therefore, most of the findings and interpretations given in the following chapters have already been published elsewhere. The focus of this paper was mainly to collect data from different working groups and combine them into a general description of the late Paleozoic glacial deposition, the onset of Permian rifting, and the associated history of lacustrine deposition which initiated the onset of the Karoo Supercycle in southern and eastern Africa. Continental glaciation in Late Carboniferous to Early Permian times occurred at numerous locations on the African continent and other former Gondwanan continents. In Africa, glacial and postglacial deposition represents the onset of the Karoo depositional cycle, which continued until the Early Triassic. In South Africa the glacial deposits are referred to as the Dwyka Formation, which has been studied intensively by numerous authors (Rust 1975; Martin, 1981; Visser, 1989); in Tanzania the succession was described by Wopfner and Kreuser (1986), Kreuser (1987), and Wopfner and Diekmann (1996). In East Africa the glaciogene sediments are best exposed in southwestern Tanzania, where the early Karoo deposits are exposed along the uplifted rift shoulders of Lake Malawi and Lake Rukwa. This paper summarizes the lithological and depositional history of a number of Tanzanian basins: the Ruhuhu, Songwe Kiwira, and Galula Basins, with respect to their glacial and postglacial depositional history. Special emphasis is placed on the accumulation of organic matter in large periglacial and postglacial lakes and the thermal history of these organic source rocks. Additionally, a stratigraphic approach is presented to compare the Tanzanian sections with other southern and eastern African localities of comparable age and focus on the ubiquitous onset of glacial and postglacial deposits in the region. The glacial nature of these deposits was first established by Spence (1957). Later, Wopfner and Kreuser (1986), Kreuser (1987), Diekmann (1993), Wopfner and Diekmann (1992), and Wopfner and Diekmann (1996) described these sequences in detail and established a modern nomenclature and a depositional model. The most detailed lithologic investigation with respect to facies distribution and regional differentiation was performed by Diekmann (1993), when the formal lithostratigraphic terminology was established. Geochemical analysis of the nature and

thermal history of Karoo source rocks was performed by Kreuser et al. (1988), Dypvik et al. (1990), Diekmann (1993), and Kreuser (1995b). The present paper summarizes the most important features and incorporates a model of the depositional environment and thermal history of these potential source rocks for hydrocarbon generation. Within the Karoo succession, several phases of lacustrine development occurred in which organic matter accumulated (Kreuser, 1995b); however, only the periglacial to postglacial development is highlighted in this paper. Similar sedimentary successions were recorded from neighboring countries in Africa: South Africa (Rust, 1975), Madagascar (Besairie, 1972), Congo (Boutakoff, 1948), Ethiopia (Worku and Astin, 1992), and Oman (Qidwai, 1988), which are mentioned here for comparative purposes. The Tanzanian succession serves as a reference section, which is comparatively well described and analyzed. STRATIGRAPHY The southwestern Tanzanian Karoo basins comprise diverse facies successions of glacial and postglacial deposits (Fig. 1): 1. Ruhuhu Basin (subdivided into a western Mchuchuma sub-basin and an eastern Ngaka sub-basin) along the eastern fault scarp of the Malawi Rift. Post-Karoo movements have dissected the Ruhuhu Basin into several half grabens, i.e., the Lumecha sub-basin. Owing to the southeastern tilt of the half grabens, outcrops of the Idusi Formation lie mainly on the northern and western margins of these blocks. 2. Songwe-Kiwira Basin, situated northwest of the Ruhuhu Basin. 3. Galula Basin, along the southern boundary of the Rukwa Rift (Dypvik et al., 1990; Mliga, 1994). The type section of the glaciogenic sequence in Tanzania is in the Ruhuhu Basin in the gorge of the Idusi River in the northwestern Mchuchuma sub-basin and is known as the Idusi Formation (Kreuser et al., 1990; Fig. 2). Comparable stratotypes were detected along the Ketewaka River in the northeastern Mchuchuma sub-basin and along the Nyamangami River of the western Ngaka sub-basin (Fig. 2). Additional sections of the Idusi Formation were obtained from logs of exploration boreholes drilled by CDC (Colonial Development Corp.) and MADINI (Geological Survey of Tanzania) and were described by Wopfner and Diekmann (1996).

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AFRICA

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Tanzania Lake Victoria

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Figure 1. Regional distribution of continental Karoo (Permian-Triassic) rocks in Tanzania and neighboring countries.

At the type locality in the Idusi River gorge the Idusi Formation reaches a thickness of 240 m. Because of facies variations, Diekmann (1993) subdivided the Idusi Formation into two lithostratigraphic units, a lower Lisimba Member and an upper Lilangu Member. Lisimba Member This member is named after a tributary of the Idusi River (thickness, 170 m, Fig. 3). The lower succession consists of different lithotypes of diamictite, conglomerate, and sandstone as well as silty mudstone, exhibiting large dropstones and lonestones (Diekmann, 1996). Higher in the succession, laminites are dominant in a mudstone matrix with abundant lonestones. The Lisimba Member exhibits a typical olive green color caused by abundant chlorite and has an arkosic composition. The thickness ranges up to 420 m owing to a strong paleorelief of the Pre-

cambrian basement (Kreuser, 1987). No palynological evidence has been found in the lower Lisimba Member, and owing to the stratigraphic position of the younger Lilangu Member, the lower member may have been deposited in Late Carboniferous times (Diekmann, 1996). The basal unit of the Lisimba Member consists of massive diamictite, grading into faintly bedded diamictite. Higher in the section, siltstone with dropstones up to 60 cm in diameter is present. The middle portion of the section consists of siltstone and mudstone with slumping structures. The upper part consists of olive green laminites with a few dropstones and sandstones. Lilangu Member Named after a tributary of the Idusi River (thickness, 70 m, Fig. 3), the Lilangu Member developed gradually from the glaciogenic succession beneath and is separated locally by an

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Figure 2. Distribution of sub-basins and geology of glacial and postglacial deposits in the Ruhuhu Basin. Note the fault-controlled basin margins along the northeast trending blocks.

unconformity from the overlying Mchuchuma Formation (Semkiwa, 1992). The boundary between the Lisimba and Lilangu Members is characterized by the first black fissile shale that is highly pyritic, containing calcareous concretions (Wopfner and Diekmann, 1996). Above this are black organic-rich siltstones, locally bituminous, intercalated with sandstone lobes and breccias. The amount of organic matter is high (kerogen types III and IV, from 3% to 6%). Palynological analyses indicate assem-

blages assigned to the Granulatisporites confluens Oppel zone, indicating a late Asselian to early Sakmarian age (Wopfner and Kreuser, 1986). Regional Lithological Variations The reference section along the Ketewaka River of the Idusi Formation measures ~100 m in thickness and exhibits faceted

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Figure 3. Composite log of the Idusi Formation at the type locality in the Idusi gorge and the northwestern Mchuchuma sub-basin. Legend is attached (for location, see Fig. 2). Based on Wopfner and Diekmann (1996).

and striated diamictites at the base and continues in a similar way along the Idusi River into laminites farther upward (Fig. 4). The Lilangu Member, however, exhibits a different facies development. Although the dominant black, organic-rich shale is present, intercalations with sandy lobes, several of which attain 10 m in thickness, are present, with an overall increased thickness of 120 m. Diekmann (1993) termed these sediments the Muhimbi facies, named after a tributary of the Ketewaka River. In the Ngaka sub-basin (Fig. 1) the “long section” of the Idusi Formation reaches a thickness of >700 m. In contrast to correlative lithologic successions, the lower 140 m is characterized

by sandstone with cross-bedding, plus siltstone and conglomerate that contain cobble-size lonestones. Diekmann (1993) termed this section the Ndongosi facies after the escarpment it forms, and interpreted it as glaciofluvial in origin. This is overlain by 160 m of massive deltaic sandstones, referred to as the Nanderuka facies (Diekmann, 1993). The Lilangu Member reaches a maximum thickness of ~300 m; common intercalations of mudstone are present, and a generally higher sand to shale ratio is observed. Both south and north of the “long section” the Idusi Formation is characterized by a considerably condensed succession with a thickness of up to 40 m. For that reason, other authors have used the term “short section” (McKinlay, 1954). The threefold lithologic subdivision and most of the described lithofacies are recognizable (Fig. 5). However, a section north of the Mkapa River exhibits a massive basal diamictite with faceted and striated clasts of various diameters resting on basement, interpreted as a lodgment till (Diekmann, 1996). It is followed by reworked diamictite and green bedded sandstones. The Lilangu Member has the typical black shales at the base, but the upper part is usually absent, probably having been removed by erosion prior to deposition of the Mchuchuma Formation (Diekmann, 1996). Songwe Kiwira Basin Approximately 50 km west of the northern end of Lake Malawi (Fig. 1), another small basin appears, named after the

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Kreuser and Woldu paleovalley filled with basal diamictites interfingering with sandstone and conglomerate, capped by a few meters of laminites. Galula Basin This small basin, along the southwestern flank of the Rukwa Rift (Fig. 1), is an erosional remnant of Karoo strata. Here, Karoo strata are preserved along the depressed side of tilted fault blocks. Mbede (1993) noted seismic surveys indicating that in the deepest part of the Rukwa graben, some 7 km below ground level, 5 km of Karoo strata is present. There is only one outcrop of the Idusi Formation along the Chizi River, of ~13 m thickness (Wopfner and Diekmann, 1996). The succession begins with a basal diamictite, overlain by laminites of the Kipololo facies, topped by laminated mudstones with microclasts. The Lilangu Member is not developed here. Mhukuru Basin Located 65 km south-southwest of Songea near the Mozambique border (Fig. 1), this basin contains a well-developed succession of the Idusi Formation, with a maximum thickness of ~120 m. The Lisimba Member exhibits basal diamictites with dropstones and sandstones and green laminites. The overlying Lilangu Member consists of 20 m of black shales (Diekmann, 1993). DEPOSITIONAL ENVIRONMENT Lisimba Member

Figure 4. Log of the Ketewaka section in the northwestern Mchuchuma sub-basin. See Figure 3 for legend. Based on Wopfner and Diekmann (1996).

main rivers Kiwira and Songwe, which cross it. The southern part of the basin continues into Malawi, exhibiting an originally larger areal extent of Karoo deposits. The northern part of the basin is covered by Quaternary basalts of the Rungwe volcanics. The Idusi Formation is restricted to parts of the Lisimba Member, starting with a basal diamictite with cobbles that locally reach 30 cm in diameter. This unit is interpreted as a lodgment tillite (Dypvik et al., 1990). Facies variations can be studied in borehole sections drilled by the Coalfield Exploration Team of the Peoples Republic of China (Anonymous, 1979). Wopfner and Diekmann (1996) presented a depositional model that shows a 60-m-deep

The basal diamictites consist of massive to faintly bedded, poorly sorted mixtures of clasts resting in a sandy-silty matrix. Clasts, derived from local basement rocks, are faceted and locally striated, which is characteristic for glacial transport (Wopfner and Kreuser, 1986). In this paper, diamictites are interpreted as subaqueous melt-out and flow tillites, intercalated with meltwater deposits. Lodgment tills rarely occur. Depositional evidence for this interpretation is given in detail in Wopfner and Kreuser (1986). In the long section they are replaced by an upwardly fining succession, starting with conglomerate and ending with mudstone. This is interpreted as a transition from braided-river to subaquatic conditions in a proglacial environment (Diekmann, 1993; Fig. 6). The dominant lithology of the Lisimba Member is characterized by silty and sandy mudstones containing rare lonestones with diameters up to 80 cm. The lonestones are interpreted as ice-rafted dropstones. The fine clastics were settling from meltwater plumes as suspended particles. A few ripple laminations and sandy layers are current indicators that were formed during subaqueous flows (Diekmann, 1996). Laminites consist of alternating silt-clay couplets (8–15 cm) with common deformation structures, and some microclasts and rare dropstones are also present, which were deposited as periglacial varves during seasonal changes from a thermally stratified

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Figure 5. Correlation chart of three sections of the Idusi Formation in the Ruhuhu Basin, displaying a typical threefold subdivision. In addition to the depositional environments, the successive climatic changes are seen at the right. Modified from Diekmann (1993).

water column (Diekmann, 1993). Diekmann interpreted the laminites as sandy layers from proximal positions that developed during meltwater discharge during times of episodically slightly higher temperatures (Fig. 6). Lilangu Member No sharp boundary could be discerned between the Lisimba and Lilangu Members. A transitional zone exists where the silt gradually disappears and is replaced by black to olive gray shale with a high TOC (total organic carbon) content of up to 0.5%. Early diagenetic concretions are commonly observed wherein pyrite replaces plant remains. These basal shales represent a deepwater lacustrine environment during a postglacial stage (Diekmann, 1993). Overlying the basal shales are black, organic-rich siltstones, reflecting a high amount of terrestrial organic debris that consists of kerogen types III and IV. TOC values of up to 12% were

detected. These siltstones contain fossil-wood remains and remnants of Gangamopteris. Rhythmites are intercalated with the massive siltstones and consist of couplets of silt and medium sand laminae. This succession is interpreted as sapropel-rich deposits of a deep-lake environment with restricted circulation in a thermally stratified water column (Diekmann, 1993). Rare dropstones probably signal seasonal ice rafts on the lake, with common graded bedding. Some lenticular sandstone bodies are intercalated into the black siltstones, which exhibit sole marks (flute and groove casts) and current lineations. These sandstones are interpreted as turbidites accumulating at the apex of subaquatic channels on delta fronts (Diekmann, 1993; Fig. 6). Toward the top of the member an increase in sand-size grains was detected, which developed parallel with a decrease in TOC. Wopfner and Diekmann (1996) interpreted these silty sandstones as part of a shallow lake environment, which alternated with upper-flow-regime sandstone deposits accompanied by lag deposits with common mud clasts.

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Figure 6. Paleogeographic model of the depositional environment during deglaciation time. (A) Periglacial setting with ice rafts (1), valley glaciers (2), glacial outwash (3), reworked glaciofluvial residue (4). (B) Postglacial lake model with laminites, rhythmites, lakeshore flora, and associated marginal deltaic lobes.

PALEOGEOGRAPHY Records of the Permo-Carboniferous glaciation in Africa are in most cases incomplete, and commonly only the last glacial advance was spared from erosion. Deeply incised paleovalleys around the centers of glaciation are known from several localities, i.e., from South Africa (Visser, 1989, 1997), Namibia (Martin, 1981), and Congo (Boutakoff, 1948). From other Gondwanan continents, comparable features have been recorded—e.g., Campos (1994) from Brazil, and Woolfe (1994) from Antarctica. In Tanzania a number of such tectonically controlled paleovalleys have preserved much of the sedimentary record, the most impressive of these being the Ruhuhu Basin, with almost 900 m of sediments. This is the only region where such a thick succession was preserved in Tanzania, attributed probably to thermal subsidence during the early phase of rifting in the Early Permian of East Africa. The Idusi Formation appears to record the last glacial advance and the succeeding stages of glacial retreat and degla-

ciation. This climatic amelioration is evinced in the lithologic development of the Idusi Formation irrespective of local facies or thickness variations. The correlation between lithology and depositional environment is shown in Figure 5. The Lisimba Member, characterized by massive diamictite, records the last glacial advance and subsequent retreat. These are end-glacial or periglacial deposits that are overlain by a proximal glaciolacustrine succession intercalated with glaciofluvial and marginal deltaic deposits. The middle part of the Lisimba Member is characterized by mudstones with dropstones that record a distal facies of a proglacial lake environment. The final deglaciation is characterized by the deposition of fine-grained, organic-rich sediments of the Lilangu Member, which took place in a large, shallow lake that developed as an expansion of formerly proglacial to periglacial lake systems. Plant remains and sporadic desiccation cracks exhibit times of exposure when the lake was reduced in extent and depth. This large lake not only filled the paleorelief but overstepped onto higher areas where it was directly in contact with Precambrian

Formation of euxinic lakes, East Africa basement. During successive deglaciation and exposure of the lowland areas, which were increasingly vegetated, deposition switched to plant debris and silty-clayey sediments that culminated in the accumulation of organic carbon–rich black shale. These deposits took place in an anaerobic environment at the lake bottom, interchanging at marginal positions with deltaic lobes and subaqueous fans (Fig. 5). The top of the Lilangu Member records the filling of the lake and subsequent alluvial fan deposition. Wopfner (1996) interpreted the angular unconformity between the Idusi Formation and the overlying Mchuchuma Formation (Mpera Sandstone Member) as a basinwide tectonic event, which in the Galula and Songwe-Kiwira Basins led to erosion and removal of the sediments of the Lilangu Member. REGIONAL TECTONIC CONTROL AND APPROXIMATE DATING The deposition of black shales during the euxinic stages of lake development of the Ruhuhu Basin and other early Karoo basins was a significant event that can be traced regionally across East Africa. The abundance of organic matter led to anaerobic conditions in those regionally extensive deposits. Alternatively, anaerobic conditions may have led to extensive preservation of organic matter. Postglacial conditions were characterized by a climatic amelioration and represented a transition from cold and arid to temperate and humid conditions. Shortly thereafter, coalrich cyclothems prevailed, which recorded seasonal variations during deposition of the Mchuchuma Formation (Kreuser, 1991). Coal deposition commenced as an almost simultaneous event throughout entire Gondwanaland, which likely contributed to a lowering of atmospheric CO2 (Wopfner and Diekmann, 1996). The unconformity between the underlying Lilangu Member and the overlying Mpera Sandstone is evidence for a basinwide tectonic event that probably led to the complete erosion of the black shale in the Galula and Songwe-Kiwira Basins. The Mpera Sandstone is interpreted as braided stream deposits that initiated the second depositional sequence of the Karoo succession (Kreuser et al., 1990). This second unit, the Mchuchuma Formation (Fig. 5), was dominated by coal swamps and is also important for the source rock assessment of the Karoo succession in Tanzania (Kreuser et al., 1988). On a regional scale in Africa south of the Sahara, glaciogenic and postglacial deposits are known from South Africa, Zimbabwe, Zambia, Malawi, Kenya, Madagascar, Congo, and Mozambique. The basal diamictite is present in Zimbabwe (Mid-Zambezi and Sabi-Lunde Basins; Kreuser et al., 1990). In Zambia glacial deposits are known from the Gwembe, Luano, Luangwa, and Barotse Basins, respectively (Kreuser et al., 1990). In Malawi there is some evidence for the glacial character of the basal beds. In Kenya some indicators point to a glacial origin of the basal beds (Martin, 1981), and a similar record was noted for Mozambique (Kreuser, 1995b). There is no doubt that a glacial character exists in Madagascar in the Morondava Basin, whereas

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an erosional hiatus separates Precambrian to Permian strata in the Majunga Basin (Besairie, 1972). For the Congo, basal diamictites are recorded and interpreted as being of glacial origin (Boutakoff, 1948). The black shale above the diamictite is not of such a ubiquitous distribution. In Zambia the Mukumba Member shows similarities to the Lilangu Member as being part of the Luwumba Formation in the Luangwa Valley, both separated by an unconformity (Kreuser et al., 1990). In Malawi the black shales are present in the Nkana and Livingstonia Basins, separated from the underlying tillites by an unconformity (Kreuser et al., 1990). Also in Madagascar, black shales follow the tillites in the Morondava Basin (Besairie, 1972). In the Congo the tillites are overlain by two separate black shale zones that locally interfinger with fluvial sandstones, both separated by local unconformities (Kreuser, 1995a). The Lilangu Member is partly developed in those basins which survived later tectonic exhumation. Locally it was not deposited at all, but deposition evolved from a glacial to a peat swamp environment, marked by either a rapid transition or a regional unconformity between the glacial and postglacial units. However, the viability of much of the data presented in the literature is not convincing, as much of the evidence has not been verified by detailed lithological descriptions; thus some of the stratigraphic nomenclature is not up to date with lithostratigraphic correlations. Additionally, dating by palynology or paleontology is poor for these continental deposits, and much more work is needed in order to establish a reliable litho-chronological framework. Apparently the Karoo basins farther east in the vicinity of the coastal area were never reached by the Paleozoic glaciation, or at least are not exposed at the surface (Kreuser et al., 1990). No basal diamictites and succeeding euxinic black shales are known from the coastal basins of Tanzania, Kenya, and Somalia (Mbede, 1997). Wopfner and Diekmann (1996) suggested a duration for the deposition of the Lilangu Member of ~5–6 m.y. In comparison with the approximate duration of the glaciation, which lasted ~20–25 m.y., deglaciation would have taken one quarter of the overall time. The revised estimate is based on palynological analyses that indicate a late Asselian to early Sakmarian age (Granulatisporites confluens Oppel zone; Wopfner and Kreuser, 1986) for the Lilangu Member of the Idusi Formation. ORGANIC GEOCHEMISTRY OF EUXINIC LACUSTRINE BLACK SHALES OF THE LILANGU MEMBER Several authors analyzed samples with high organic contents from Lower Permian lacustrine strata from East Africa in order to estimate their hydrocarbon potential (Kreuser et al., 1988; Dypvik et al., 1990; Kagya, 1991; Diekmann, 1993). Only a few samples were collected from the postglacial deposits, however; most of the samples came from coal seams in the Mchuchuma Formation or from Upper Permian deposits in rift lake environments (Mpanju,

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1999). For the sake of comparison, some of these results are mentioned here as well, but our focus is on the black shale of the Lilangu Member. The hydrocarbon potential of the Lilangu Member black shale was evaluated by Rock-Eval pyrolyses, and the hydrogen index (HI in mg HC/g TOC) was measured. We also analyzed the Tmax (°C) values during successive heating of samples, and this was matched with vitrinite reflection data. Temperature ranges (Rock-Eval pyrolyses) for most of the samples were from 425 °C to 460 °C, with the black shale from the Lilangu Member showing the highest temperatures. These samples exhibit a fair hydrocarbon generation potential, ranging from 65 to 150 mg HC/g TOC. Some TOC values reach 11%, but most are <5% (Fig. 7A). The kerogen of the dispersed organic matter in the black shale is dominantly type III. Vitrinite reflection values of the black shale average 0.97%, which is a little lower than expected from pyrolysis data (Kreuser, 1995b). Macerals are dominated by inertinite and vitrinite (Semkiwa et al., 2003). In comparison, coals of the Mchuchuma Formation show HI values up to 300 mg HC/g TOC with slightly lower Tmax values (Fig. 7B). The generation potential from the black shale is mainly gas-prone, depending on the maceral composition. However, lamalginite (the most oil-prone liptinite maceral) and exsudatinite (remobilized bitumen) indicate the onset of bituminization and liberation of mobile hydrocarbons, both present in the black shale below the coals (Semkiwa et al., 2003). Sediments and coals from the Ilima colliery in the SongweKiwira Basin and in the Rukwa Basin, measured by Dypvik et al. (1990), show fair HI values, although they exhibit high variations

(0–224 mg HC/g TOC). Tmax is in the range of 430 °C to 458 °C (Fig. 7C). Unfortunately, nothing is known about the maceral composition and maturity derived from these vitrinite reflection measurements. No specific information about the stratigraphic position is available, so it is not known if the black shale was sampled or is the age equivalent of the Lilangu Member in the Ruhuhu Basin. Also Mpanju (1999) analyzed the hydrocarbon potential of the Karoo coals in the southwestern rifts of Tanzania, and his results are generally in accordance with the analyses described above. Kreuser (1995b) calculated a subsidence history for the Ruhuhu Basin using vitrinite reflection, palynological dating, and lithostratigraphic correlation with an inferred geothermal gradient of 25 °C km–1. Maximum subsidence of the basal glacial units of 2.6 km was attained during the Early Triassic. Afterward there might have been a slight uplift of several hundred meters during the breakup unconformity of Gondwanaland in Middle Jurassic time. A strong uplift occurred only during the late Miocene, which brought basal rocks to the surface along the uplifted basin margins. Several other Karoo rifts in Tanzania, and probably in neighboring areas, were exposed to temperatures conducive to the formation of light hydrocarbons. This has not been proven for many rifts, however, and it might be possible that lower geothermal gradients and subsidence values for the rift centers were more favorable locally for the generation of liquid hydrocarbons. In turn, these hydrocarbons might have migrated into overlying porous rock that acted as suitable reservoirs and/or traps for hydrocarbon accumulation. More research is needed to shed more light on this matter.

Figure 7. Composite HI (hydrogen index) versus Tmax (maximum temperature) diagrams for interpretation of kerogen type and maturity level of organic-rich source rocks. (A) Lilangu Member of the Ruhuhu Basin in Tanzania (modified from Diekmann, 1993). (B) Sediments and coals from the Permian Mchuchuma Formation. Open circles, Ketewaka sub-basin; full circles, Ngaka sub-basin; full squares, Mchuchuma sub-basin. (C) Sediments and coals from the Songwe-Kiwira and Rukwa Basins (adapted from Dypvik et al., 1990). TOC—total organic carbon.

Formation of euxinic lakes, East Africa CONCLUSIONS The glacial and postglacial deposits of Tanzania and other countries of eastern and southern Africa show a remarkable homogeneity in lithology and distribution. The lower part of the sequence was deposited in many basins of the region, mainly in incised paleovalleys, and later was exposed owing to uplift and tilting during Tertiary rift development. The twofold subdivision into a lower diamictite (Lisimba Member) and an upper lacustrine euxinic postglacial unit (Lilangu Member) has been well established in Tanzania and can be correlated with many neighboring successions in the region. During early Permian times, a high TOC content in the black shale of the Lilangu Member records a euxinic lake environment in a postglacial setting in most of southern and eastern Africa. These black shales represent source rocks for the generation of natural gas and some liquid oil and are within a favorable temperature window. Even though they probably were deposited ubiquitously in most of the Permian basins of the region, distribution was also controlled by the regional and local tectonic framework. The transition from a euxinic to lacustrine environment of a cool climate to a moderately warm climate, where coal swamps developed, is often disguised by an erosional hiatus. In combination with the overlying coal deposits, these organic-rich deposits represent a potential source rock for the entire East and South African region, which could be of some future economic potential. ACKNOWLEDGMENTS The authors are indebted to the sponsoring institutions of Deutsche Forschungsgemeinschaft and Deutscher Akademischer Austauschdienst and are grateful for the collaboration with the former working group members Heli Wopfner, Berni Diekmann, Pascal Semkiwa, Charles Kaaya, and Stefan Markwort. The first author is grateful for the convener of International Geological Congress 34 in Florence, Italy, Oscar López-Gamundí, for the invitation to publish in this GSA Special Paper. REFERENCES CITED Anonymous, 1979, Coalfield Geological Exploration Team of the P.R. of China, 1979, report on the geological exploration of the Songwe-Kiwira coalfield: Beijing, United Republic of Tanzania, unpublished report, 223 p. Besairie, H., 1972, Géologie de Madagascar. I. Les terrains sédimentaires: Annales Géologiques de Madagascar, Fascule 35, Tananarive. Boutakoff, N., 1948, Les Formations Glacières et Post-Glacières Fossilifères d’Age Carbonifère de la Région Walikale (Kivu, Congo Belge): Louvain, Belgium, Institut Géologie Université Louvain, Memoire 9, 124 p. Campos, J.E., 1994, A glaciaçao Neopalozoica na porçao meridional da Bacia Sanfranciscana: Revista Brasileira de Geociencias, v. 24, p. 65–76. Diekmann, B., 1993, Paläoklima und glazigene Karoo-Sedimente des späten Paläozoikums in SW-Tansania: Köln, Geologisches Institut Universität Köln, Sonderveröffentlichungen, 90, 193 p. Diekmann, B., 1996, Petrographic and diagenetic signatures of climatic changes in peri- to post-glacial Karoo sediments of SW-Tanzania: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 125, p. 5–25, doi: 10.1016/0031 -0182(95)00084-4.

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Dypvik, H., Nesteby, H., Ruden, F., Aagard, P., Johansson, T., Msindai, J., and Massay, C., 1990, Upper Paleozoic and Mesozoic sedimentation in the Rukwa-Tukuyu region, Tanzania: Journal of African Earth Sciences, v. 11, p. 437–456, doi: 10.1016/0899-5362(90)90022-7. Kagya, M.J., 1991, The source rock potential of the Nyasa rift basin, oil shows in Tanzania: Journal of Southeast Asian Earth Sciences, v. 5, p. 407–419, doi: 10.1016/0743-9547(91)90055-3. Kreuser, T., 1987, Late Paleozoic glacial sediments and transition to coal bearing Lower Permian in Tanzania: Fazies, v. 17, p. 149–158. Kreuser, T., 1991, Facies evolution and cyclicity of alluvial coal deposits in the Lower Permian of East Africa (Tanzania): Geologische Rundschau, v. 80, p. 19–48, doi: 10.1007/BF01828766. Kreuser, T., 1995a, Tectonics and climatic controls of lacustrine sedimentation in pre-rift and rift settings in the Permian-Triassic of East Africa: Journal of Paleolimnology, v. 13, p. 3–19, doi: 10.1007/BF00678108. Kreuser, T., 1995b, Rift to drift evolution in Permian-Jurassic basins of East Africa, in Lambiase, J., ed., Hydrocarbon Habitat in Rift Basins: Geological Society [London] Special Publication 80, p. 297–315. Kreuser, T., Schramedei, R., and Rullkötter, J., 1988, Gas prone source rocks from cratogene Karoo basins in Tanzania: Journal of Petroleum Geology, v. 11, p. 169–184, doi: 10.1111/j.1747-5457.1988.tb00811.x. Kreuser, T., Wopfner, H., Kaaya, C.Z., Markwort, S., Semkiwa, P.M., and Aslanidis, P., 1990, Depositional evolution of Permo-Triassic basins in Tanzania with reference to the economic potential: Journal of African Earth Sciences, v. 10, p. 151–167, doi: 10.1016/0899-5362(90)90052-G. Martin, H., 1981, The Late Paleozoic Gondwana glaciation: Geologische Rundschau, v. 70, p. 480–496, doi: 10.1007/BF01822128. Mbede, E., 1993, Tectonic Development of the Rukwa Rift Basin in SW Tanzania: Berliner geowissenschaftliche Abhandlungen, Reihe A, 152, 92 p. Mbede, E., 1997, The coastal basins of Somalia, Kenya and Tanzania, in Selley, R.C., ed., Sedimentary Basins of the World, v. 3: Amsterdam, Elsevier, p. 211–233. McKinlay, A.C.M., 1954, Geology of the Ketewaka-Mchuchuma Coalfield, Njombe District: Geological Survey of Tanganyika Bulletin 21, 46 p. Mliga, N.R., 1994, Depositional environment, stratigraphy and hydrocarbon potential of the Rukwa rift basin, SW Tanzania [Ph.D. thesis]: Durham, North Carolina, Duke University, 172 p. Mpanju, F.J., 1999, Hydrocarbon potential of Karoo coals and associated rocks of the SW rift basins of Tanzania: Journal of African Earth Sciences, v. 28, p. 52–53. Qidwai, H.A., 1988, Evidence of Permian–Carboniferous glaciation in the basal Murbat Sandstone Formation, southern region, Sultanate of Oman: Journal of Petroleum Geology, v. 11, p. 81–88, doi: 10.1111/j.1747-5457.1988 .tb00802.x. Rust, I.C., 1975, Tectonic and sedimentary framework of Gondwana basins in southern Africa, in Campbell, C.S.W., ed., Gondwana Geology: Symposium, 3rd, Canberra, Australia, p. 537–564. Semkiwa, P.M., 1992, Depositional Environment and Coal Petrography of Permian Coal Deposits in Karoo Basins of SW Tanzania: Köln, Geologisches Institut der Universität Köln, Sonderveröffentlichungen 84, 184 p. Semkiwa, P.M., Kalkreuth, W., Utting, J., Mpanju, F., and Hagemann, H., 2003, The geology, petrology, palynology and geochemistry of Permian coal basins in Tanzania: 2. Songwe-Kiwira Coalfield: International Journal of Coal Geology, v. 55, p. 157–186, doi: 10.1016/S0166-5162(03)00108-3. Spence, J., 1957, The Geology of Part of the Eastern Province of Tanganyika: Geological Survey of Tanganyika Bulletin 84, 126 p. Visser, J.N.J., 1989, The Permo-Carboniferous Dwyka Formation of Southern Africa: Deposition by a predominantly subpolar marine ice sheet: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 70, p. 377–391, doi: 10.1016/0031-0182(89)90115-6. Visser, J.N.J., 1997, A review of the Permo-Carboniferous glaciation in Africa, in Martini, I.P., ed., Late Glacial and Postglacial Environmental Changes, Quaternary, Carboniferous-Permian and Proterozoic: Oxford, UK, Oxford University Press, p. 169–191. Wopfner, H., 1996, The late Paleozoic Idusi Formation of SW Tanzania. Record of change from glacial to post-glacial conditions: Journal of African Earth Sciences, v. 22, p. 575–595, doi: 10.1016/0899-5362(96)00038-3. Wopfner, H., and Diekmann, B., 1992, Neue Ergebnisse aus der spätpaläozoischen glazigenen Abfolge in der Karoo Tansanias, Zentralblatt Geologie Paläontologie: Teil 1, p. 2689–2700. Wopfner, H., and Diekmann, B., 1996, The Late Paleozoic Idusi Formation of SW Tanzania: A record of change from glacial to postglacial conditions:

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Worku, T., and Astin, T.R., 1992, The Karoo sediments (Late Paleozoic to Early Jurassic) of the Ogaden Basin, Ethiopia: Sedimentary Geology, v. 76, p. 7–21, doi: 10.1016/0037-0738(92)90136-F.

MANUSCRIPT ACCEPTED BY THE SOCIETY 21 DECEMBER 2009

Printed in the USA

The Geological Society of America Special Paper 468 2010

Stratigraphic and paleofloristic record of the Lower Permian postglacial succession in the southern Brazilian Paraná Basin Roberto Iannuzzi* Paulo A. Souza* Michael Holz* Departamento de Paleontologia e Estratigrafia, Instituto de Geociências, Universidade Federal do Rio Grande do Sul, Cx.P. 15.001, CEP. 91.501-970, Porto Alegre, RS, Brazil

ABSTRACT A correlation between plant zones and palynozones within a sequencestratigraphic framework for the upper Paleozoic rocks of the Paraná Basin, Brazil, is attempted for the first time. Our study indicates that (1) the stratigraphic boundaries (lithostratigraphic boundaries and sequence boundaries) are not coincident with most of the biostratigraphic horizons; (2) the boundaries between palynozones and plant zones are also not coincident with each other; and (3) the boundaries of palynozones lie near the maximum flooding surfaces through the interval analyzed. The emerging pattern suggests that plant zones are mostly controlled by depositional processes and palynozones by climate-driven changes. Therefore, the plant zones correspond to distinct ecofacies, and are better regarded as ecozones rather than as biozones. On the other hand, the climatic changes that affected the palynofloras were related directly to the most significant transgressive events, suggesting a link with eustatic oscillations caused by Early Permian Glacial-Interglacial Phases. Aims of further work may include (a) evaluation of taphonomic controls in plant-bearing beds, (b) better understanding of the relation between the plant-bearing strata and their equivalent palynologic zones, (c) correlation between palynologic and paleobotanic data and the sequence-stratigraphic framework already established in other areas, and (d) improvement of the chronostratigraphic chart of the basin through the discovery of layers suitable for radiometric dating.

INTRODUCTION Stratigraphic and paleontologic studies of sedimentary successions from the southernmost Brazilian Paraná Basin, comprising deposits in the states of Santa Catarina and Rio Grande do Sul, have been carried out by many researchers since the begin-

ning of the twentieth century. Most of these studies were motivated by the presence of economically important coal reserves in the area. In the last decades, increased knowledge on the stratigraphy and paleontology of the deposits from Rio Grande do Sul resulted in the establishment of sequence-stratigraphic and biostratigraphic frameworks applicable to the southernmost part of

*E-mails: Iannuzzi—[email protected]; Souza—[email protected]; Holz—[email protected]. Iannuzzi, R., Souza, P.A., and Holz, M., 2010, Stratigraphic and paleofloristic record of the Lower Permian postglacial succession in the southern Brazilian Paraná Basin, in López-Gamundí, O.R., and Buatois, L.A., eds., Late Paleozoic Glacial Events and Postglacial Transgressions in Gondwana: Geological Society of America Special Paper 468, p. 113–132, doi: 10.1130/2010.2468(05). For permission to copy, contact [email protected]. ©2010 The Geological Society of America. All rights reserved.

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the Paraná Basin (Guerra-Sommer and Cazzulo-Klepzig, 1993; Holz, 1999; Souza and Marques-Toigo, 2003). However, no attempts have been made to integrate sequence-stratigraphic and biostratigraphic data. This paper is intended as a first approach to integrating the stratigraphic sequences and biozones described in the sections from the Rio Grande do Sul State. It is part of a major biostratigraphic study, including the revaluation of the biozonations already proposed. The main objective of this study is to understand the stratigraphic significance and external controls on these biozones and to evaluate the associated paleoecologic factors in order to establish a broad biostratigraphic framework for the Lower Permian deposits of the Paraná Basin in the future. GEOLOGIC SETTING The southern margin of the Paraná Basin, an Ordovician to Cretaceous intracratonic basin in the South American platform, is located in southernmost Brazil and north/northwestern Uruguay (Fig. 1A). The basin covers a surface area ~1,700,000 km2, has a NE-SW elongated shape, and is ~1750 km long and 900 km wide. The sedimentary fill of the basin was controlled by tectonic-eustatic cycles linked to late Paleozoic orogenic events caused by subduction and terrain accretion on the southwestern margin of the Gondwana continent, and ended with the opening of the South Atlantic during the Mesozoic (e.g., Zalán et al., 1990; Milani et al., 1994). The prevalence of eustatic-tectonic cycles, which controlled sedimentation in Paraná Basin, has generated a stratigraphic record that is marked by numerous interruptions caused by erosion and non-deposition. Milani et al. (1994) subdivided basin fill into six second-order depositional sequences, referred to as “Supersequences”: Rio Ivaí (Rio Ivaí Group, Ordovician–Silurian), Paraná (Paraná Group, Devonian), Gondwana I (Tubarão and Passa Dois Groups, Carboniferous– Permian), Gondwana II (several Triassic formations), Gondwana III (São Bento Group, Jurassic-Cretaceous), and Bauru (Bauru Group, Cretaceous). Figures 1A and 1B show the distribution of these strata within the southernmost Brazilian areas of the Paraná Basin (Rio Grande do Sul and Santa Catarina States) as well as the upper Paleozoic basin stratigraphy. The stratigraphic interval discussed in this paper comprises Lower Permian rocks (ranging in age from Asselian/Sakmarian to late Artinskian/early Kungurian), encompassing the following lithostratigraphic units (in stratigraphic order): Itararé Group, Rio Bonito Formation, Palermo Formation, and Irati Formation (Fig. 1B). This succession constitutes the lower portion of the Gondwana I Supersequence of Milani et al. (1994), which forms the thickest sedimentary sequence of the basin (up to 2300 m). The lowermost section of this second-order sequence occurs only in the depocenter of the basin (from north Santa Catarina to the Mato Grosso State), assigned to the Late Carboniferous and corresponding to the basal and middle portions of the Itararé Group (see Petri and Souza, 1993, and Souza, 2006). From the Early Permian onwards, glacially related strata onlap basin margins, as in Rio Grande do Sul, recording a second-order transgressive

cycle that began with the deposition of the upper Itararé unit and has its maximum flooding surface within the Palermo Formation (e.g., Milani et al., 1994; Holz, 1999). EARLY PERMIAN SUCCESSION IN THE SOUTHERN PARANÁ BASIN The lowermost unit of the Early Permian succession in Rio Grande do Sul, the Itararé Group, was formed in pro- and postglacial environments. The Rio Bonito Formation is composed of fluvio-deltaic and estuarine to shallow-marine deposits, and the uppermost Palermo Formation has a shallow-marine (lower shoreface to offshore) origin. The Rio Bonito Formation is fluvial and deltaic at the base, turns into estuarine in the middle portion, and is coastal to shallow marine in the upper part. Details may be seen in Holz (2003) and Holz et al. (2006). Detailed (third-order) sequence-stratigraphic analysis of this stratigraphic interval within the study area, carried out by our group (e.g., Holz, 1999, 2003; Holz et al., 2000, 2006), revealed that important regional base-level changes generated three distinct regional unconformities, which constitute sequence boundaries roughly corresponding to the lithostratigraphic limits (Fig. 2). The lowermost limit is SB 1, an unconformity regionally marked by the contact between the Early Permian succession and Neoproterozoic crystalline rocks of the basement (or locally metasedimentary rocks of the early Paleozoic Camaquã Basin), enclosing a hiatus up to 300 Ma. SB 2 is marked by the contact between fluvial and deltaic sandstone belonging to the Rio Bonito Formation and the underlying glacio-marine mudstone of the Itararé Group. This noticeable facies shift is indicative of an unconformity, marking a relative sea-level drop which is also recorded at other locations around the Rio Grande do Sul shield (e.g., Alves and Ade, 1996; Holz, 1997). SB 3 is marked by an erosive surface generated by fluvial incisions on underlying estuarine and shallow-marine mudstone and sandstone (within the upper interval of the Rio Bonito Formation). These unconformities define three third-order depositional sequences, which were mapped throughout the study area on the basis of subsurface data, using the maximum flooding surface of the second sequence as a datum. AGE CONSTRAINTS IN THE PERMIAN OF THE PARANÁ BASIN The main biostratigraphic problem of most of Gondwanan Carboniferous-Permian deposits, including those in the Paraná Basin, is the absence of biostratigraphically significant marine faunal elements (such as foraminifers or ammonoids), which prevent correlations with the Late Carboniferous and Early Permian international stages. In addition, radiometric data are still scarce and conflictive in the Permian sequences of Gondwana, militating against accurate age calibration of the available palynostratigraphic schemes. Thus, correlation with international stratigraphic stages can be considered as difficult and speculative for the most

A SAN

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Geologic Key Cenozoic coverage Mesozoic rocks (Paraná Basin) Paleozoic rocks (Paraná Basin) Paleoproterozoic basement

1

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Studied locations:

1 - Morro do Papaleo outcrop 2 - Quiteria outcrop 3 - Borehole CA-53

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B Passa Dois Group

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Serra Alta Formation: shallow marine

Corumbataí Formation: tidal plain

Tubarão Group

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Palermo Formation: shallow marine

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Dourados Formation: continental

Rio Bonito Formation: shore face, deltaic, lagoonal, fluvial

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Itararé Subgbroup: glacio-marine to glacio-continental

V

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Irati Formation: marine restricted

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Teresina Formation: shoreface

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GONDWANA I SUPERSEQUENCE (Upper Carboniferous - Permian)

Rio do Rasto Formation: deltaic, lagoonal fluvial, eolian

. .. ........ V .. .. ..... .. V .. .. .. .. V . .... V V

Aquidauana Formation: glacio-continental

.... .. . . . . .. .. .. .......................................... . .......................................... . . .. .. .. .. .. .. .. .. .. .. .. .. .. .. .. .. .. .. .. . . . . .................. . . . .............. ................ . . .. . ................. . . .................

.................. ................ Serra Geral Formation: ............. magmatic (Jurassic/Cretaceous) ..........

Figure 1. (A) Location and distribution of the Paraná Basin in southern Brazil (small map) and the main stratigraphic units of the Rio Grande do Sul State (large map) showing the studied outcrops. (B) Chronostratigraphic chart for the upper Paleozoic succession of the Paraná Basin with an overview on spatial relations of the deposits throughout the basin and the main depositional environments; the rectangular box refers to the section analyzed in this paper (modified from Souza, 2006).

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Chrono- Lithostratigraphy 3rd Order Stratigraphic Sta. Catarina Stratigraphy Framework

Idealized Stratigraphic Profile

SEQ-3

Artinskian PERMIAN

MFS-3

HST2 MFS-2

Coastal (delta, shoreface, barrier, lagoon, tidal flat) sandstones and shales

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Fluvial sandstones, conglomerates and silts

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Paraguassu Mb.

Rio Bonito Fm.

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Shallow-marine shales and mudstones

TST3 TS-3

Siderópolis Mb.

Triunfo Mb.

Itararé Gr. SEQ-1

LEGEND Interbedded sandstones and coal seams (deltaic, lagoonal systems)

HST3

Palermo Fm.

Rio Grande do Sul

Glacio-continental to glacio-marine sedimentites

TS-2

LST2 SB-2 SB-1

Interval not studied

Pre-Cambrian basement

Figure 2. Sequence-stratigraphic framework and lithostratigraphic equivalents of the studied interval. Lithostratigraphy is based on Schneider et al. (1974), third-order sequence stratigraphy is from Holz et al. (2000), and ages are based on the palynomorph biostratigraphy provided by Souza and Marques-Toigo (2003, 2005) and the chronostratigraphic scheme established after Césari (2007), Iannuzzi (2007), Stephenson (2008), and Guerra-Sommer et al. (2008a, 2008b). Abbreviations: SEQ—third-order sequences, SB—sequence boundary, TS—transgressive surface, MFS—maximum flooding surface, LST—lowstand systems tract, TST—transgressive systems tract, HST—highstand systems tract.

part. For all these reasons, an additional stratigraphic control based on extra-basinal correlations with Namibian (Aranos Basin) and South African (Main Karoo Basin) sequences is proposed herein to support a more consistent chronostratigraphic framework for the Lower Permian strata of the Paraná Basin. In the Itararé Group of Rio Grande do Sul, the lowermost strata are positioned slightly above the stratigraphic levels containing marine shells related to “Eurydesma Fauna,” which are typically recorded toward the central parts of the basin (RochaCampos and Rösler, 1978). In the Main Karoo Basin of South Africa, deposits associated with the “Eurydesma Fauna” encompass a time span of ~10 m.y, based on under- and overlying tuffs dated through radiometric methods by Bangert et al. (1999). Hence, the “Eurydesma Fauna” ranges from the earliest Asselian (ages of 302 ± 3 Ma and 299.2 ± 3.2 Ma from Bangert et al., 1999) up to the middle Sakmarian stage (288 ± 3 Ma and 289.6 ± 3.8 Ma according to Bangert et al., 1999). This fauna has been used as an important stratigraphic marker throughout Gondwanan sections, being associated with a widespread transgressive event referred to as the “Eurydesma transgression” by Dickins (1984). Thus, this independent calibration suggests that the stratigraphic interval studied begins somewhere in the early to midSakmarian (Fig. 2). Recently, de Matos et al. (2001) obtained a date of 267.1 ± 3.4 Ma (U/Pb) from a tonstein interbedded in the lower coal seam of Candiota Coal, positioned in the middle Rio Bonito Formation, from Rio Grande do Sul, which corresponds to the Protohaploxypinus goraiensis Subzone of the Vittatina costabi-

liz Zone (according to Souza and Marques-Toigo, 2005). In the context of the time scale of Gradstein et al. (2004), this age is assigned to the Roadian Stage. However, this same coal seam and the other one above were recently redated by Guerra-Sommer et al. (2008b), giving new ages that varied from 288.4 ± 2.6 to 293 ± 3.5 Ma for the four interbedded tonsteins analyzed. These ages correspond to earliest to mid-Sakmarian interval, according to the above-mentioned time scale. In addition, Guerra-Sommer et al. (2008a, 2008b) furnished new radiometric zircon ages of 285.4 ± 8.6 Ma and 288.76 ± 1.4 Ma geochronologic data from a tonstein interbedded with other coal seams of the Faxinal coalfield, also situated in the middle Rio Bonito Formation from Rio Grande do Sul. These dates correspond to the Hamiapollenites karroensis Subzone of Vittatina costabiliz Zone (VcZ). According to the time scale of Gradstein et al. (2004), this dating corresponds to the earliest mid- to latest Sakmarian interval. These ages have been partially contested by some authors (Césari, 2007; Iannuzzi, 2007; Stephenson, 2008) who disagree with the dating postulated by the biostratigraphic framework, which is based mostly on palynomorphs but also on plants, marine invertebrates, and terrestrial vertebrates used in the correlative sections from the other Gondwanan deposits. We have therefore discarded the Roadian age obtained by de Matos et al. (2001) and the oldest ages proposed by Guerra-Sommer et al. (2008a, 2008b). On the other hand, Guerra-Sommer et al. (2008b) established a mean average age of 290.6 ± 1.5 Ma for tonsteins of both the

Southern Brazilian Paraná Basin Candiota and Faxinal coalfields, indicating an age for this coal succession constrained to middle Sakmarian. This age interval is reasonable and matches up approximately with the ages previously established by fossil content, and we accept it herein as the middle portion of the Rio Bonito Formation (Fig. 2). In conclusion, the deposits of the Rio Bonito Formation start in the earliest middle Sakmarian and extend up to the latest Sakmarian or even the earliest Artinskian, considering the age obtained for the overlying Irati Formation and the relatively precise dating of this unit. Finally, the Irati Formation, the uppermost unit studied herein, was considered as Late Artinskian in age (according to Gradstein et al.’s time scale), based on the dating of 278.4 ± 2.2 Ma obtained from tuff beds through the SHRIMP zircon method by Santos et al. (2006). Presence of the “Mesosaurus Fauna” and of equivalent lithofacies indicates that this unit is correlated with the Whitehill Formation. In Namibia (Aranos Basin) and South Africa (Main Karoo Basin), the age of the Whitehill Formation is given by a U/Pb age of 270 ± 1 Ma (= Kungurian-Roadian boundary according to the time scale of Gradstein et al., 2004) as determined from tuffs in the Collingham Formation of the Main Karoo Basin in South Africa (Turner in Stollhofen et al., 2000). This unit overlies the Whitehill Formation, indicating a probable Artinskian/ Kungurian age for the Whitehill deposits. This dating agrees with the late Artinskian age obtained from tuffs in the Irati Formation of the Paraná Basin (Santos et al., 2006). Consequently, the top of the stratigraphic interval under consideration is neither older than Late Artinskian nor younger than Kungurian (Fig. 2). PALEONTOLOGIC SETTING The most complete palynologic and floral zonations for the upper Paleozoic strata of the Paraná Basin have been proposed by Daemon and Quadros (1970) and Rösler (1978), respectively. Although these schemes are based on extensive sampling, they are informal and lack detail. Accordingly, these zonations were abandoned in the last decades in favor of more detailed schemes. The following formal palynozones and plant zones proposed are restricted to the southernmost Paraná Basin. Palynostratigraphy Marques-Toigo (1991) formally proposed the first palynostratigraphic scheme for the upper Paleozoic in the southern portion of the Paraná Basin. Further refinements were published by Souza and Marques-Toigo (2003, 2005). Permian palynozones of the southernmost Paraná Basin were established by Souza and Marques-Toigo (2005) and correspond to interval biozones, in accordance with the formal criteria of the International Subcommission on Stratigraphic Classification (ISSC) of the International Union of Geological Sciences (IUGS) (Murphy and Salvador, 1999). Selected taxa of these palynozones, including common and diagnostic species previously illustrated by Iannuzzi and Souza (2005), are shown in Figure 3.

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Vittatina costabilis Interval Zone Bilaterally and radially symmetrical monosaccate pollen grains are the most common palynomorphs of this palynozone, such as Cannanoropollis, Plicatipollenites, Caheniasaccites, and Potonieisporites, constituting up to 50%–60% of assemblages, as well as cingulizonate spores (e.g., Lundbladispora, Cristatisporites). Disaccate pollen grains became dominant only within the upper portion of this palynozone, considered to be correlated with the Hamiapollenites karroensis Subzone. Common taxa are Limitisporites, Vittatina, Scheuringipollenites, and Protohaploxypinus. Spores are common and only locally dominant in the coal beds representing local parautochthonous/autochthonous flora (Punctatisporites, Horriditriletes, Lundbladispora, and Cristatisporites), where they represent up to 80% of the assemblages. The first appearance of polyplicate species of the genus Vittatina (V. saccata, V. subsaccata, V. costabiliz, V. vittifera), as well as species of Protohaploxypinus (P. goraiensis, P. limpidus), Fusacolpites fusus and Illinites unicus, marks the lower limit of this zone, recognized within the upper Itararé Group. The upper limit is marked by the appearance of diagnostic species of the Lueckisporites virkkiae Interval Zone, within the uppermost Rio Bonito Formation and/or the lowermost Palermo Formation. The Vittatina costabiliz Zone is divided into two units, the Protohaploxypinus goraiensis and Hamiapollenites karroensis Subzones (in ascending stratigraphic order). The former is characterized by the range of Protohaploxypinus goraiensis, Illinites unicus, and Protohaploxypinus limpidus. The latter is defined by the range of the eponymous species; its base corresponds to the first appearance of Striatopodocarpites fusus and Staurosaccites cordubensis. This interval zone correlates with the H3-J intervals of Daemon and Quadros (1970) and with the Cannanoropollis korbaensis Zone of Marques-Toigo (1991). Lueckisporites virkkiae Interval Zone Taeniate and polyplicate pollen grains, mainly those of the genus Protohaploxypinus, Striatopodocarpites, Striatoabieites, Lunatisporites, Lueckisporites, Vittatina, Weylandites, and Marsupipollenites, are dominant in this zone. Common monosaccate and disaccate pollen grains (Plicatipollenites, Potonieisporites, Limitisporites) of the underlying zone are less common. In general, spores are scarce; among them monolete species are more common (Thymospora, Laeviagatosporites). Two new species appear within this zone (Thymospora criciumensis and Convolutispora pintoi). Lueckisporites virkkiae seems to have a stratigraphically consistent first appearance in the Paraná Basin, at the base of the K interval of Daemon and Quadros (1970), which is considered a datum throughout the Paraná Basin. The last appearance of Hamiapollenites karroensis and Stellapollenites talchirensis, and the first appearance of Lueckisporites virkkiae, L. densus, L. stenotaeniatus, Pakhapites fasciolatus, Weylandites lucifer, Protohaploxypinus hartii, P. sewardi, P. microcorpus, Lunatisporites variesectus, Alisporites nuthallensis, and Striatopodocarpites pantii characterize the lower limit of this zone in the uppermost Rio Bonito Formation and the

A

D

G

J

B

C

E

F

H

K

I

L

Figure 3. Selected pollen taxa of the Vittatina costabiliz Interval Zone (based on Souza and Marques-Toigo, 2005, and Iannuzzi and Souza, 2005). (A) Protohaploxypinus goraiensis (Potonié and Lele) Hart 1964 (slide MP-P: 324, England Finder coordinate: R35/1). (B) Illinites unicus Kosanke emend. Jansonius and Hills 1976 (MP-P: 4033, U50). (C) Vittatina costabiliz Wilson 1962 (MP-P: 3573, G36/2). (D) Striatopodocarpites fusus (Balme and Hennelly) Potonié 1958 (MP-P: 312, U27/2). (E) Hamiapollenites karroensis (Hart) Hart 1964 (MP-P: 1534, Q35/2). (F) Fusacolpites fusus Bose and Kar 1966 (MP-P: 333/287, P44). Selected pollen taxa of the Lueckisporites virkkiae Interval Zone (based on Souza and Marques-Toigo, 2005, and Iannuzzi and Souza, 2005). (G) Thymospora criciumensis Quadros, Marques-Toigo and Cazzulo-Klepzig 1995 (MP-P: 1447, J24). (H) Lueckisporites densus Cauduro 1970 (MP-P: 21, N35-4). (I) Vittatina subsaccata Samoilovich 1963 (MP-P: 2541, L40/2). (J) Marsupipollenites striatus (Balme and Hennelly) Foster 1975 (MP-P: 4033, Q45). (K) Lunatisporites variesectus Archangelsky and Gamerro 1979 (MP-P: 2344, M49). (L) Staurosaccites cordubensis Archangelsky and Gamerro 1979 (MP-P: 2345, M32/2). Slides are housed at the Departamento de Paleontologia e Estratigrafia/UFRGS (under codes “MP-P”). Scale bar—20 µm.

Southern Brazilian Paraná Basin lowermost Palermo Formation. This interval zone is correlated with the K+L intervals of Daemon and Quadros (1970).

119

Guerra-Sommer and Cazzulo-Klepzig (1993) proposed a preliminary plant stratigraphic zonation containing two assemblage zones geographically restricted to Rio Grande do Sul. A new plant stratigraphic scheme was recently proposed by Jasper et al. (2003), apparently as a replacement for the above-mentioned scheme. However, this scheme was not established in accordance to the criteria and recommendations indicated by the ISSC (as summarized by Murphy and Salvador, 1999), because it lacks (a) information on co-occurring diagnostic taxa, (b) stratotypes, (c) a reference section, and (d) stratigraphic ranges of taxa. In addition, no correlation between previous and recent schemes was established. Therefore, this scheme is discarded herein. The two plant zones correspond to assemblage biozones as established by Guerra-Sommer and Cazzulo-Klepzig (1993) based on a few outcrops. All these localities were referred as “stratotypes” to these zones by Guerra-Sommer and CazzuloKlepzig (1993, p. 64 and 66). However, they are better understood as type localities containing type sections. In reality, no specific stratotype has been proposed for each zone/subzone. Apart from this, these plant zones are in accordance with the criteria of the International Subcommission on Stratigraphic Classification of IUGS, as generally applied to paleobotanic units (Cleal, 1999). Selected taxa are illustrated in Figures 4 and 5, including main diagnostic taxa and accessory elements. Recently, Iannuzzi et al. (2003a, 2003b, 2006), Jasper et al. (2003, 2005), Tybusch (2005), Souza and Iannuzzi (2007, 2009), Vieira et al. (2007), and Tybusch and Iannuzzi (2008) have included new key taxa, and extended the range of some taxa previously known in these zones. The present scheme is outlined below and shown in Figure 6.

nate whereas Glossopteris-like leaves are generally rare and represented by few species (Glossopteris indica, G. communis). The protoglossopterid elements (e.g., Rubidgea lanceolata, R. obovata), initially described by Guerra-Sommer and CazzuloKlepzig (1993) for this zone, have been discarded by Tybusch and Iannuzzi (2008). The latter authors considered all Rubidgealike leaves as different types of Gangamopteris ones. Sphenopsid stems (Paracalamites-like) and leaf branches, glossopterid foliage and platispermic seeds are abundant, whereas pteridophylls (Botrychiopsis plantiana) and cordaitalean foliage (Cordaites hislopi = C. spathulata) are common, and ginkgoalean leaves (Chiropteris sp.), conifer-like leaf shoots, and glossopterid fructifications are relatively rare. The lower limit of this zone is defined by the first appearance of glossopterid elements in the upper Itararé Group, whereas the upper limit is mostly marked by the appearance of diagnostic species of the overlying Glossopteris/Rhodeopteridium Zone in the middle-upper Rio Bonito Formation (Fig. 6). This plant zone was originally divided into two units, the Gangamopteris obovata and Phyllotheca indica Subzones. However, the species Phyllotheca indica, eponym of the last subzone, has been regarded by some authors as a synonym of Phyllotheca australis (in Rohn and Lages, 2000). In fact, considering the sphenopsid material used by Guerra-Sommer and Cazzulo-Klepzig (1993) to erect the P. indica Subzone and the information provided by these authors, it seems reasonable to conclude that Brazilian specimens attributed to P. indica must be included in P. australis. Consequently, the denomination P. indica Subzone is herein formally replaced by the P. australis Subzone. The Gangamopteris obovata Subzone seems to be defined only by the ranges of Cornucarpus patagonius and Cordaicarpus truncata (Iannuzzi et al., 2003a, 2003b, 2006; Jasper et al., 2003; Souza and Iannuzzi, 2007, 2009; Tybusch and Iannuzzi, 2008). Therefore, this subzone is mainly defined by the absence of the diagnostic species of the overlying Phyllotheca australis Subzone (Fig. 6). The Phyllotheca australis Subzone is a more consistent unit, defined by the first appearance of the eponymous species as well as by the following taxa: Stephanophyllites cf. S. sanpaulensis, Cheirophyllum speculare, Kawizophyllum sp., Glossopteris occidentalis, Samaropsis kurtzii, and S. gigas (Fig. 6). The Gangamopteris obovata Subzone was originally recognized in the upper Itararé Group, and the Phyllotheca australis Subzone has been placed in the lowermost Rio Bonito Formation (Guerra-Sommer and Cazzulo-Klepzig, 1993). However, Iannuzzi et al. (2003a, 2003b, 2006) suggested that the range of the Phyllotheca indica Subzone be restricted to the top of the Itararé Group.

Botrychiopsis plantiana Assemblage Zone This zone is characterized by the first appearance of glossopterid elements and by the local abundance of the pteridophyll Botrychiopsis plantiana and/or Phyllotheca-like sphenophytes. Among the glossopterid plants, the gangamopterid elements (Gangamopteris buriadica, G. angustifolia, G. obovata) domi-

Glossopteris/Rhodeopteridium Assemblage Zone In this zone, the genus Glossopteris becomes dominant with the emergence of new species (G. mosesii, G. browniana), displacing Gangamopteris as the most common glossopterid element. In addition, lycophytes (Brasilodendron pedroanum) and true ferns (Pecopteris sp., Asterotheca sp., Sphenopteris sp.,

Reference Wells The stratigraphic distribution proposed by Marques-Toigo (1991) and Souza and Marques-Toigo (2005) was mainly based on coal mine wells, as follows: 2-TG-69-RS well (drilled in the locality of Santa Terezinha Coal), 5-CA-03-RS, 5-CA-41-RS (Charqueadas Coal), 5-CA-91-RS (Gravataí-Morungava Coal), 2-TG-88-RS (Chico Lomã Coal), P7 (Iruí Coal), and N3 (Santa Rita Coal). Geologic and palynologic data from these wells were published by Marques-Toigo and Pons (1974), Bortoluzzi et al. (1980), Marques-Toigo et al. (1982, 1984), and Picarelli et al. (1987). Plant Biostratigraphy

A

D

B

E

F

C

G

Figure 4. Selected plant taxa of the Botrychiopsis plantiana Assemblage Zone. (A) Gangamopteris buriadica. (B) Gangamopteris obovata. (C) Cordaites sp. (D) Phyllotheca indica (= P. australis). (E) Cheirophyllum speculare. (F) Botrychiopsis plantiana. (G) Stephanophyllites cf. S. sanpaulensis. Specimens housed at the Departamento de Paleontologia e Estratigrafia, Universidade Federal do Rio Grande do Sul.

A

B

C

D

F

E

G H

Figure 5. Selected plant taxa of the Glossopteris/Rhodeopteridium Assemblage Zone. (A) Glossopteris occidentalis. (B) Botrychiopsis valida. (C) Asterotheca sp. (D) Coricladus quiterensis. (E) Rhodeopteridium sp. (F) Lycopodites sp. (G) Leaves of Brasilodendron sp. (H) Stem of Brasilodendon pedroanum. Specimens housed at the Departamento de Paleontologia e Estratigrafia, Universidade Federal do Rio Grande do Sul.

122

Iannuzzi et al. GEOCHRONOLOGY LITHOSTRATIGRAPHY

SAKMARIAN

GUATÁ GROUP RIO BONITO FM.

ITARARÉ GROUP BOTRYCHIOPSIS PLANTIANA ZONE

BIOSTRATIGRAPHY

GANGAMOPTERIS OBOVATA SUBZONE

PHYLLOTHECA INDICA SUBZONE

Cornucarpus patagonicus Cordaicarpus truncata * Gangamopteris obovata Botrychiopsis plantiana § Chiropteris spp. * Gangamopteris angustifolia § Gangamopteris buriadica § Glossopteris indica § Glossopteris communis Cordaites hislopi Buriadia isophylla § Samaropsis seixasi Cordaicarpus aff. C. famatinensis * Stephanophyllites spp. * Cheirophyllum speculare * Dicranophyllum sp. * Samaropsis kurtzii * Samaropsis cf. S. rigbyi * Samaropsis gigas * Cordaicarpus cerronegrensis * Cordaicarpus aff. C. brasilianus * Phyllotheca australis Kawizophyllum sp. * Glossopteris occidentalis § Scutum sp. Rhodeopteridium sp. Sphenopteris spp. * Pecopteris pedrasica * Neomariopteris sp. * Asterotheca sp. * Brasilodendron pedroanum Cyclodendron sp. * Ginkgophytopsis sp. * Botrychiopsis valida * Glossopteris mosesii Glossopteris browniana G. obovata var. major * Plumsteadia sennes Arberia minasica Otokaria sp. Coricladus quiteriensis * Samaropsis aff. S. millaniana *

Neomariopteris sp.) appear for the first time in Rio Grande do Sul. The appearance of a new conifer-like plant (Coricladus quiteriensis) and Gangamopteris leaves (G. obovata var. major) is another significant feature. Consequently, the plant assemblages become highly diversified. Assemblages of this zone correspond

/ARTINSKIAN

GLOSSOPTERIS / RHODEOPTERIDIUM ZONE

Figure 6. Updated biostratigraphic distribution of plant taxa within the assemblage biozones of Guerra-Sommer and Cazzulo-Klepzig (1993) based on new data sets published by Iannuzzi et al. (2003a, 2003b, 2006), Jasper et al. (2003, 2005), Souza and Iannuzzi (2007, 2009), Vieira et al. (2007), and Tybusch and Iannuzzi (2008). Taxa marked with “*” represent new key species added recently. Taxa marked with “§” represent previous taxa whose stratigraphic range has been modified based on the new data set. Ages of biozones are based on the palynostratigraphy provided by Souza and Marques-Toigo (2003, 2005) and the chronostratigraphic scheme established after Césari (2007), Iannuzzi (2007), Stephenson (2008), and GuerraSommer et al. (2008a, 2008b).

to the typical “Glossopteris Flora” and typically occur in association with coal measures. The lower limit of this zone is apparently defined by the last record of Stephanophyllites cf. S. sanpaulensis and Cheirophyllum speculare, and by the first appearance of Brasilodendron

Southern Brazilian Paraná Basin pedroanum, Pecopteris spp., Asterotheca sp., Neomariopteris sp., Sphenopteris spp., Botrychiopsis valida, Glossopteris browniana, G. mosesii, Arberia minasica, Ottokaria sp., Plumsteadia sennes, and Coricladus quiteriensis within the middle-upper Rio Bonito Formation (Fig. 6). Its upper limit is marked by the disappearance of the most diagnostic species in the overlying Palermo Formation. According to available data, the B. valida Subzone of Jasper et al. (2003) and the Glossopteris/Rhodeopteridium Zone of Guerra-Sommer and Cazzulo-Klepzig (1993) are considered equivalent units for stratigraphic purposes. Thus, the new taxa introduced by Jasper et al. (2003) as belonging to the B. valida Subzone were transferred to and listed in the Glossopteris/ Rhodeopteridium Zone for the present analysis (Fig. 6). Reference Sections The reference sections for the Botrychiopsis plantiana Zone are in the localities of Morro do Papaléo (including both Gangamopteris obovata and Phyllotheca australis Subzones) and Quitéria (the upper subzone only) (Guerra-Sommer and CazzuloKlepzig, 1993). Additional localities were listed for the lower subzone, G. obovata, including Faxinal, Goulart Farm, Acampamento Velho, and Cambaí Grande. The reference sections for the Glossopteris/Rhodeopteridium Zone occur in Quitéria and Faxinal Mine (Guerra-Sommer and Cazzulo-Klepzig, 1993). In addition, the uppermost part of the section from Morro do Papaléo is regarded as a stratotype for the Glossopteris/Rhodeopteridium Zone (Iannuzzi et al., 2003a, 2003b). Consequently, we chose herein the localities of Morro do Papaléo and Quitéria for a detailed lithologic logging of plantbearing exposures and stratigraphic studies, due to the presence of more than one of the biostratigraphic units erected by GuerraSommer and Cazzulo-Klepzig (1993) (see Piccoli et al., 1991, and Iannuzzi et al., 2003a, 2003b, 2006). In Morro do Papaléo, all three successive biostratigraphic units previously established have been recorded, whereas in Quitéria at least the two uppermost units (Phyllotheca australis Subzone and Glossopteris/ Rhodeopteridium Zone) are present. DISCUSSION Integration of Biostratigraphic and SequenceStratigraphic Information Figures 7, 8, and 9 show the results obtained from the integration of stratigraphic and paleontologic data. Palynozones are more precisely correlated with the stratigraphic scheme of Holz (1999) than plant zones because palynozones are based on well samples. The Vittatina costabiliz Interval Zone occurs from the base of Sequence 1, extending through the upper Itararé Group and most of the Rio Bonito Formation. Its upper limit lies near the maximum flooding surface of Sequence 3, encompassing the uppermost Rio Bonito Formation and the lowermost Palermo Formation (Fig. 9). The Lueckisporites virkkiae Interval Zone

123

occurs from the middle interval of Sequence 3 to at least the top of Sequence 4 in the Irati Formation (Fig. 9). The boundary between the Protohaploxypinus goraiensis and the Hamiapollenites karroensis Subzones lies near the maximum flooding surface MFS 2, but its exact position cannot be traced with the available data. The stratigraphic position of plant zones is indicated by correlations between the Morro do Papaléo and Quitéria outcrops and the CA-53 borehole (Fig. 7) (Holz, 1999). This correlation was established on the basis of outcrop data. The Morro do Papaléo section is the most complete, containing the three plant zones (Fig. 8). The Botrychiopsis plantiana Zone extends throughout Sequence 1, being its upper limit coincident with SB 2, which corresponds to the Itararé/Rio Bonito boundary (Fig. 9). The precise position of the boundary between the Gangamopteris obovata and Phyllotheca australis Subzones remains unknown. The Glossopteris/Rhodeopteridium Zone passes through SB 3, extending approximately from SB 2 to TS 3 in the upper Rio Bonito Formation (Fig. 9). With respect to correlations between the plant zones and palynozones, the Botrychiopsis plantiana Zone is equivalent to the Protohaploxypinus goraiensis Subzone of the Vittatina costabiliz Interval Zone (Fig. 8). The Glossopteris/Rhodeopteridium Zone corresponds mostly to the Hamiapollenites karroensis Subzone, being normally recorded in the upper Rio Bonito Formation. Nevertheless, the presence of lycopsid stems (assigned with doubts to the genus Cyclodendron) in coal seams of the Morro Papaléo (Fig. 8), and of Brasilodendron pedroanum in coal seams of the Candiota Mine (Chaloner et al., 1979), indicates that the Glossopteris/Rhodeopteridium Zone extends well into the Protohaploxypinus goraiensis Subzone (Fig. 9). In contrast to the original proposal of Guerra-Sommer and Cazzulo-Klepzig (1993), the presence of the Phyllotheca australis Subzone in the Quitéria outcrop has not been confirmed herein. Despite the occurrence of some diagnostic taxa related to this subzone in the lowermost strata of the Quitéria exposure, such as Botrychiopsis plantiana and Chiropteris sp., the palynologic content indicates a higher stratigraphic position for this outcrop. Among the palynomophs recovered, an important index element of the Hamiapollenites karroensis Subzone is Striatopodocarpites fusus (Jasper et al., 2006). This indicates that the Quitéria exposure corresponds to the uppermost Vittatina costabiliz Interval Zone, being stratigraphically close to the Rio Bonito/Palermo boundary (Figs. 8 and 9). The presence of typical elements of the P. australis Subzone in Quitéria outcrop therefore should be interpreted as a result of a facies-controlled reappearance of those taxa at a higher stratigraphic position. Consequently, the Glossopteris/Rhodeopteridium Zone is considered the only biozone present in the Quitéria outcrop (Fig. 8), and the above-mentioned taxa correspond to co-occurring elements present in the lower levels, within the underlying B. plantiana Zone. Another important aspect that emerges from this analysis is the fact that the lithostratigraphic boundaries are not coincident

LST 2

TST 2

HST 2

Distal borehole (CA - 53)

?

GLACIAL CONTINENTAL GLACIAL MARINE

FLUVIO-DELTAIC

LAGOONAL

FLUVIO-DELTAIC

LAGOONAL MFS 2

SB 3

Approx. 30 km

2

SB 1

SB

-C- C- C-

-C- C- C-

-C- C- C-

-C- C- C-

FLUVIO-DELTAIC

QUITÉRIA OUTCROP

FLUVIAL

FAN DELTAIC

Approx. 20 km

FLUVIAL

PALYNOZONES

MORRO DO PAPALÉO OUTCROP

Towards Inner Valley - Continent

Figure 7. Stratigraphic correlation between outcrops analyzed herein and borehole CA-53. Palynologic zonation (right side) was very useful for determining the stratigraphic position of the Quitéria section. Transect made along a paleovalley (“Mariana Pimentel Paleovalley”), being the distal facies represented by the well and the proximal ones by outcrops. Sequence stratigraphy shown at left side from Holz et al. (2000), palynostratigraphy at right side according to Souza and Marques-Toigo (2003, 2005). Stratigraphic abbreviations, see Figure 2; palynozones, see the text.

SEQUENCE STRATIGRAPHY

H. karroensis Sub. P. goraiensis Sub.

Towards Platform - Shoreline

Vittatina costabilis Zone

H. karroensis Subzone

-C- C- C-

-C- C- C-

-C- C- C-

-C- C- C-

V. costabilis I. Zone

Glossopteris / Rhodeopteridium A. Zone

a, d

b

n, o p, q

Figure 8. Stratigraphic range of plant fossils: Quitéria and Morro do Papaléo (based on Piccoli et al., 1991; Iannuzzi et al., 2003a, 2003b, 2006; Jasper et al., 2003, 2005). Palynozones and plant zones are plotted side by side for each outcrop. Taxa: a—Botrychiopsis plantiana; b—Gangamopteris angustifolia; c—Gangamopteris obovata; d—Chiropteris sp.; e— Phyllotheca indica (= P. australis); f—Stephanophyllites sanpaulensis; g—Cheirophyllum speculare; h—?Dicranophyllum sp.; i—Gangamopteris buriadica; j—Kawizophyllum sp.; k—Glossopteris indica; l—Glossopteris occidentalis; m— Cyclodendron sp.; n—Gondwanostachis sp.; o—Arberia sp.; p—Glossopteris browniana; q—Rhodeopteridium sp.; r— Brasilodendron cf. pedroanum; s—Coricladus quiteriensis; t—Lycopodites sp.; u—Ginkgophytopsis sp.; v—Asterotheca sp.; w—Pecopteris sp.; x—Sphenopteris sp.; y—Neomariopteris sp. Abbreviations: SB—erosive surface/sequence boundary; B. plantiana A. Zone—Botrychiopsis plantiana Assemblage Zone; Glossopteris/Rhodeopteridium A. Zone— Glossopteris/Rhodeopteridium Assemblage Zone; V. costabiliz I. Zone—Vittatina costabiliz Interval Zone; G. obovata Sub— Gangamopteris obovata Subzone; P. australis Sub—Phyllotheca australis Subzone.

Quitéria

0m

1m

r-t

SB 2

0m

5m

a

b

c-i

j-l

m

r, u

n, v - y

Morro do Papaléo

B. plantiana A. Zone

SB 3

Vittatina costabilis Interval Zone

No record

Glossopteris / Rhodepteridium A. Zone G.obovata Sub.. P.australis Sub.

No record Protohaploxypinus goraiensis Subzone

126

Iannuzzi et al.

Stages

Ages (Ma)

Plants

Stratigraphy

Palynomorphs

Kungurian 275.6 + 0.7

Lithologies

Main Systems

Sea level

+

??????????

Irati

278.4 + 2.2

Sequences

Environment

No plant zones

Lueckisporites virkkiae Interval Zone

Formation

Shallow carbonatic marine

-

Marine

Period Epoch

Biostratigraphy

Sequence 4

Chronostratigraphy

Artinskian

B. plantiana Assemblage Z.

No record

Sequence 3

MFS 3

Fluvial / Delta plain

Coastal Plain

Estuarine / Lagoonal

Sequence 2

Rio Bonito Formation

Shoreline

MFS 2

Shoreline

(upper) Itararé Group

SB 2 Shoreline Prodelta Glaciomarine

Pro- Postglacial glacial

Sakmarian

Shallow marine

SB 3

Sequence 1

288.4 + 2.6

H. karrooensis Subzone Protohaploxypinus goraiensis Subzone

285.4 + 8.6

Glossopteris / Rhodeopteridium Assemblage Zone

284.4 + 0.7

Vittatina costabilis Interval Zone

Palermo Fm.

Cisuralian

P E R M I A N

SB 4

Figure 9. Correlation chart of chrono- and biostratigraphic (= based on palynomorphs and plants) data, and paleoenvironmental framework for the Lower Permian of the Paraná Basin in Rio Grande do Sul State. Note the relations among lithostratigraphic units, stratigraphic sequences, biozone boundaries, paleoenvironments, and sea-level changes. Differences between palynozones and plant zones seem to be strongly facies controlled. Chronostratigraphy according to Gradstein et al. (2004). Abbreviations: SB—sequence boundary, MFS—maximum flooding surface. Legend: Levels with three stars indicate radiometric ages obtained by Guerra-Sommer et al. (2008a, 2008b) from tonsteins of the Rio Bonito Formation and by Santos et al. (2006) from tuff beds of the Irati Formation. See text for more details.

with most biostratigraphic horizons. The Vittatina costabiliz Interval Zone includes the upper Itararé Group and most of the Rio Bonito Formation (Fig. 9). No significant biostratigraphic difference has been recorded within these sections, despite noticeable lithologic changes below and above SB 2, mapped regionally throughout the state of Rio Grande do Sul. The boundary between the V. costabiliz and Lueckisporites virkkiae Interval Zones is recorded in the uppermost Rio Bonito Formation, slightly below the Rio Bonito/Palermo boundary (Fig. 9). The only coincidence is that of the Botrychiopsis plantiana Zone and Sequence 1 (Fig. 9). With respect to the boundaries between palynologic zones and plant stages, the same situation is noted. The transition from the Botrychiopsis plantiana Zone to the Glossopteris/ Rhodeopteridium Zone occurs within the V. costabiliz Interval Zone and is not coincident with the Protohaploxypinus goraiensis/ Hamiapollenites karroensis boundary (Fig. 9).

Correlations of the Biozones Palynozones can apparently be extended throughout the Paraná Basin, taking into account the range of selected sporepollen grains and the horizons of their appearance and disappearance. Therefore, a correlation between these units and the previous informal proposal of Daemon and Quadros (1970) for the whole basin was established by Souza and Marques-Toigo (2003). According to these authors, the V. costabiliz Interval Zone correlates with the H3-J intervals of Daemon and Quadros (1970), and L. virkkiae Interval Zone is correlated with the K and L intervals. In terms of extra-basinal correlations, the V. costabiliz Interval Zone is correlated with the Cristatisporites Zone and the Fusacolpites fusus-Vittatina subsaccata Interval Biozone from northern Argentina (Césari and Gutiérrez, 2000). Other similar assemblages occur elsewhere in Lower Permian deposits of

Southern Brazilian Paraná Basin Gondwana regions, such as Australia, India, Antarctica, Oman, and Saudi Arabia (e.g., Jones and Truswell, 1992; Stephenson and Filatoff, 2000). Like the Vittatina costabiliz Zone, the L. virkkiae Interval Zone has equivalent units widely known from the Middle to Upper Permian Gondwanan strata (Césari and Gutiérrez, 2000). In contrast, plant zones are only valid for the Rio Grande do Sul State, as originally proposed by Guerra-Sommer and Cazzulo-Klepzig (1993). In fact, the other few formal plant zones proposed in the last decades for the Paleozoic deposits of the Paraná Basin are restricted in terms of both time intervals and geographic distribution within the basin (Millan, 1987; GuerraSommer and Cazzulo-Klepzig, 1993). For this reason, Iannuzzi and Souza (2005) recently proposed three successive informal floras, representing developmental stages of the whole flora, to be adopted for the Lower Permian strata of the Paraná Basin. A tentative correlation between these floras and plant zones can be attempted on the basis of similarities in taxonomic lists and relative abundance of floral elements. As with the other Permian floral records in Gondwana, the predominance of gangamopterid leaves over glossopterid ones and the abundance of sphenophytes are not unique to the B. plantiana Zone but are also typical features shared among the earliest Permian assemblages of the “Phyllotheca–Gangamopteris Flora” from the Paraná Basin, as defined by Iannuzzi and Souza (2005). Similar assemblages occur elsewhere in the lowermost lower Gondwana strata (= Sakmarian), as for instance in Australia, India, southern Africa, and southern South America (Retallack, 1980; Anderson and Anderson, 1985; Archangelsky et al., 1996; Maheshwari and Bajpai, 2001). Glossopteris-dominated assemblages, on the other hand, similar to those of the younger Glossopteris/Rhodeopteridium Zone, are typically in coal-bearing strata throughout Gondwana. In the Paraná Basin, they correspond to the “Glossopteris–Brasilodendron Flora” (Iannuzzi and Souza, 2005). Closely related Lower Permian (= Sakmarian–Artinskian) assemblages occur in Argentina, South Africa, India, and Australia (Retallack, 1980; Anderson and Anderson, 1985; Archangelsky et al., 1996; Maheshwari and Bajpai, 2001). An integrative correlation chart of chrono-, litho-, and biostratigraphic data for the Lower Permian of the Paraná Basin is shown in Figure 10. Significance of the Biozones General Controls According to Iannuzzi and Souza (2005), the “Phyllotheca– Gangamopteris Flora” records the transition between deglaciation and postglacial depositional sequences in the uppermost Itararé Group. This flora was therefore preserved in nearshore deposits after the retraction of the ice sheet. Coeval assemblages belonging to the Botrychiopsis plantiana Zone occur only in the uppermost Itararé Group in postglacial deposits representing proximal settings within lacustrine to lagoonal environments. The assemblages of the Glossopteris/Rhodeopteridium Zone that correspond to the “Glossopteris–Brasilodendron Flora” of

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Iannuzzi and Souza (2005) are locally abundant in the overlying Rio Bonito Formation. In Rio Grande do Sul, these assemblages occur in peat-forming and fluvial environments. Iannuzzi and Souza (2005) considered the “Glossopteris–Brasilodendron Flora” as recording the vegetation of various environments in coastal lowlands during the postglacial phase, including peatforming swamps. Plant-bearing deposits of the Rio Bonito Formation accumulated in widespread coastal plains developed immediately after the last maximum sea-level rise caused by the melting of glaciers in the Paraná Basin (Iannuzzi and Souza, 2005). Ice melt occurred during deposition of Sequence 1 (Figs. 2 and 9). As a result, the higher diversity of the Glossopteris/Rhodeopteridium Zone or “Glossopteris–Brasilodendron Flora” was previously associated with a postglacial climatic amelioration (GuerraSommer and Cazzulo-Klepzig, 1993; Jasper et al., 2003). However, the palynologic record indicates that such amelioration took place much earlier, with the appearance of assemblages of the Vittatina costabiliz Interval Zone near the base of Sequence 1 (Fig. 9). The lower limit of the V. costabiliz Zone is characterized by the appearance of different types of disaccate striate pollen (Vittatina spp. and Protohaploxypinus spp., among others), which has classically been considered as indicative of warmer conditions by palynologists from Gondwana. Despite changes in plants (from G. obovata–P. australis Subzones to Glossopteris/Rhodeopteridium Zone) and lithology (from subglacial-proglacial to paralic facies) through the uppermost Itararé Group and lower-middle Rio Bonito Formation (Figs. 2 and 9), no significant variation has been recorded within the V. costabiliz Interval Zone up to the P. goraiensis/H. karroensis boundary, except for the appearance of facies-controlled species locally restricted to coal seams (i.e., the Caheniasaccites ovatus Ecozone of Souza and Marques-Toigo, 2005). Although evidence of proglacial and subglacial environments is recorded at the base of Sequence 1 (e.g., dropstones and striated pavements corresponding to the lowermost Itararé Group in Rio Grande do Sul), the spore/pollen-producing plants are essentially the same as occur in overlying stratigraphic levels. These data are interpreted herein as reflecting the absence of significant large-scale climatic changes throughout Sequence 1 and much of Sequence 2 (Fig. 9). In conclusion, glacial evidence at the base of the interval under consideration appears to represent local rather than regional ice influence, perhaps reflecting the last ice tongues surrounded by relatively rich floras. The dominance of pollen species with conifer and pteridosperm affinities (e.g., monosaccate pollen) in the V. costabiliz Interval Zone suggests a temperate forest coexisting with the locally distributed ice tongues. Consequently, the whole interval analyzed herein is understood as having been generated during a “postglacial climatic phase,” representing the ending of the Gondwanan Ice Age, at least in the Paraná Basin. Thus, the rise of diversity in the Glossopteris/Rhodeopteridium Zone seems to have been driven by changes in the depositional settings preserved rather than by climatic changes. The characteristic elements of this zone are interpreted as colonizers of new habitats

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Palynology Chronostratigraphy

Early to Middle Permian (Cisuralian - Guadalupian)

Roadian Kungurian

Lithostratigraphy

Serra Alta / Teresina / Rio do Rasto Fms.

Daemon and Quadros (1970)

L1 L2 L3

L

Irati Fm.

Artinskian

Rio Bonito Formation

Souza and Marques-Toigo (2003, 2005)

Iannuzzi and Souza (2005)

Hamiapollenites karroensis Subzone

I2 - I4

I

I1

Vittatina costabilis Interval Zone

Protohaploxypinus goraiensis Subzone

H3

Polysolenoxylon - Glossopteris Flora

(fauna and radiometric ages)

270 + 1 Ma

No Plant Zones

Mesosaurus Fauna 278 + 1 Ma

Glossopteris - Brasilodendron Flora Phyllotheca - Gangamopteris Flora

Glossopteris / Rhodeopteridium Zone B. plantiana Zone

285.4 + 8.6 Ma 288.4 + 2.6 Ma 289.6 + 3.8 / 288.0+ 3 Ma

Eurydesma-type Fauna

Sterile Interval

Asselian

Late Carboniferous (Pennsylvanian)

Guerra-Sommer and Cazzulo-Klepzig (1993)

Sterile Interval

Sakmarian

Itararé Subgroup

Stratigraphic control

Zones from Rohn & Rösler (2000)

Lueckisporites virkkiae Interval Zone

K J

Palermo Fm.

Paleobotany

H

302 + 3.0 / 299.2+ 3.2 Ma

H1 H2

Crucisaccites monoletus Interval Zone

Westphalian G

Ahrensisporites cristatus Interval Zone

Pre-Glossopteris Flora

No Record Radiometric ages: Brazilian rocks African rocks

Figure 10. Correlation chart of chrono- and biostratigraphic (= based on palynomorphs and plants) data, and paleoenvironmental framework for the Lower Permian of the Paraná Basin and Rio Grande do Sul State with correlative radiometric dated horizons from Brazil (Irati Formation), southern Namibia (Aranos Basin), and South Africa (Main Karoo Basin). Absolute ages according to Stollhofen et al. (2000), Santos et al. (2006), and Guerra-Sommer et al. (2008a, 2008b). See text for more details. Chronostratigraphy according to Gradstein et al. (2004). Modified from Iannuzzi and Souza (2005).

associated with paralic environments recorded during deposition of Sequence 2. Climatically driven changes in palynofloras appear to have occurred only during the Vittatina costabiliz/Lueckisporites virkkiae transition. Most of the new pollen grains that arose from the Lueckisporites virkkiae Interval Zone have been considered to be derived from xerophyllic plants, based on their probable botanical affinities and intrinsic morphologies adapted to dry conditions (e.g., striate bodies). Furthermore, the co-appearance of abundant carbonate deposits within these sections indicates a general rise in average temperatures and a decrease in humidity. In addition, the analysis of growth rings of logs recovered from the Irati Formation indicates a wet-dry seasonal climate of Mediterranean type (Alves, 1999). This climatic change associated with maximum flooding surfaces present in the Palermo and Irati Formations seems to indicate the establishment of warmer climates in the Paraná Basin with increasingly dry conditions. This event may correspond to a late Artinskian/early Kungurian warm peak that followed the cold to cool temperate conditions recorded in Western Australia (Nicoll and Metcalfe, 1997). Therefore, a rise in sea level resulted from deglaciation in Eastern Gondwana

during the Artinskian-Kungurian transition (Ziegler et al., 1997). Accordingly, this flooding event could be considered as a datum throughout Gondwana, permitting an intra-continental correlation as good as the “Eurydesma Transgression” delineated by Dickins (1984) for the Early Permian (= early Sakmarian). In fact, some authors have also addressed the use of other Carboniferous– Permian postglacial transgressions as tentative chronostratigraphic horizons in Gondwana deposits (López-Gamundí, 1989; López-Gamundí et al., 1994; Veevers and Powell, 1987). Plant Zones From their analysis of assemblages from the Morro do Papaléo outcrop, Iannuzzi et al. (2003a, 2003b, 2006, 2007) suggested that the subzones included in the B. plantiana Zone are probably facies controlled. In this context, the lowermost assemblage in the Morro do Papaléo outcrop that corresponds to the G. obovata Subzone (Fig. 8) consists of transported plant material derived from nearby vegetation and preserved in shallow lake or lagoon fills. Up-section in the Morro do Papaléo, the next assemblage (which is related to the P. australis Subzone, Fig. 8) has been interpreted as riparian or marginal-lake to lagoon vegetation

Southern Brazilian Paraná Basin (Iannuzzi et al., 2003a, 2003b, 2006, 2007). Consequently, the assemblages of the B. plantiana Zone were classified by these authors as (1) allochthonous, consisting of drift plant remains associated with lower G. obovata Subzone, (2) parautochthonous, composed of short-distance transported material, and (3) autochthonous, dominated exclusively by portions of sphenophytes in growth position, with (2) and (3) being assigned to the same horizon and belonging to the upper P. australis Subzone (see Fig. 8). Furthermore, no significant sedimentologic break or environmental change was detected in the interval (5–15 m thick) that separates the fossiliferous horizons containing typical elements of the G. obovata Subzone from those of the P. australis Subzone (Fig. 8). Therefore, taking into account the Morro do Papaléo section, the plant subzones of the B. plantiana Zone seem to reflect primarily changes in the depositional settings, apparently corresponding to distinct ecofacies (Iannuzzi et al., 2003a, 2003b). Thus, the G. obovata and P. australis Subzones are better understood as ecozones rather than as biozones. During recent fieldwork at the Morro do Papaléo locality, additional recovered plant material demonstrates the presence of almost the same taxa as in the levels attributed to both the G. obovata and P. australis Subzones, except for the occurrence of B. plantiana, which still remains restricted to the lower G. obovata Subzone in this section (Fig. 8). However, the range of B. plantiana is known to extend stratigraphically upward from the base of the Quitéria outcrop, which lies within the Glossopteris/Rhodeopteridium Zone (Fig. 8). Therefore, the only exclusive taxon left in the G. obovata Subzone is Cornucarpus patagonicus, a small seed rarely found in the floral record of Rio Grande do Sul (Fig. 6). On the other hand, among all the fossiliferous localities mentioned in Rio Grande do Sul by Guerra-Sommer and Cazzulo-Klepzig (1993) (Faxinal, Goulart Farm, Acampamento Velho, Cambaí Grande, Faxinal Mine, and Quitéria), the successive and simultaneous presence of both the G. obovata and P. australis Subzones in the same exposure has been recorded only in the Morro do Papaléo section until now. Furthermore, the exact stratigraphic position of the other outcrops is not well determined yet. For all these reasons, it is not clear whether the subdivision of B. plantiana Zone could be useful for stratigraphic purposes. Given that the record of the Glossopteris/Rhodeopteridium Zone reflects distinct plant-bearing depositional settings across the coastal wetlands, at least four types of assemblages can be distinguished according to Jasper and Guerra-Sommer (1999), Jasper (2004), and Iannuzzi et al. (2003a, 2003b, 2007). Near the top of the Morro do Papaléo section, there are two distinct assemblages: (1) one is parautochthonous and consists of a sorting of distinct parts of lycophytes and glossopterid and cordaitalean plants; and (2) another is hypo-autochthonous and dominated by ferns and glossopterid foliage. Both are preserved in floodplain deposits (Fig. 8). Two additional assemblage types occur in the lower and middle intervals of the Quitéria outcrop: lower hypoparautochthonous assemblages composed of plant remains transported a short distance and preserved in lake- or lagoon-margin

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deposits, and an upper autochthonous assemblage that is dominated by lycophytes in living position, corresponding to vegetation associated with peat-forming environments (i.e., “roofshale floras”) (Fig. 8). Despite differentiation of all these assemblages, the taphonomic processes involved in the generation of these plant assemblages have not yet been adequately studied, preventing the establishment of a strong relation between microhabitats and associated floral elements. In addition, we need a better understanding of the ecologies of individual species and/or the associated communities. CONCLUDING REMARKS Until recently a proper understanding of the stratigraphic significance of established plant zones and palynozones in the Paraná Basin was obscured by the lack of an adequate stratigraphic control. Stratigraphic information from Rio Grande do Sul has accumulated during the last decades, however, resulting in the establishment of a sequence-stratigraphic framework. Thus, we have been able to attempt, for the first time, a correlation of the plant zones with palynozones and stratigraphic framework, with the following results: (1) the stratigraphic boundaries (lithostratigraphic boundaries and sequence boundaries) are not coincident with most of the biostratigraphic horizons; (2) the boundaries between palynozones and plant zones are also not coincident with each other; and (3) the boundaries of palynozones lie near the maximum flooding surfaces through the interval analyzed. These results suggest that plant zones are controlled mostly by depositional processes and palynozones by climate-driven changes. Consequently, the plant zones correspond to distinct ecofacies, and are better regarded as ecozones rather than as biozones. On the other hand, the climatic changes that affected the palynofloras were related directly to the most significant transgressive events recorded throughout the Paraná Basin. This suggests that the primary mechanism driving the changes in climatic conditions was linked to eustatic oscillations caused by Early Permian Glacial-Interglacial Phases recorded in Gondwana. The persistence of the same palynozone (i.e., V. costabiliz Zone) from the upper Itararé Group to the upper Rio Bonito Formation, despite the noticeable facies changes from pro- and subglacial to coastal-plain environments, is underscored in this study. This pattern appears to reveal the absence of significant large-scale climatic changes during this interval, because the spore/pollen-producing plants present from the bottom to the top can be considered basically the same. In addition, palynologic data suggest that the Permian glacial record of the Paraná Basin represents in reality a “postglacial phase,” equivalent to an interglacial phase. If this is the case, it is reasonable to conclude that the peak of the last glacial event (i.e., the glacial phase) did not get preserved in the sedimentary record. Thus, the plant zones proposed by Guerra-Sommer and Cazzulo-Klepzig (1993) seem to reflect primarily changes in the

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physical environment. The reappearance in the Glossopteris/ Rhodeopteridium Zone of taxa characteristic of the lower B. plantiana Zone indicates a strong facies control. In this context, the increase in floral diversity of the Glossopteris/Rhodeopteridium Zone can be directly linked with the record of new lowland environments within the Rio Bonito Formation rather than with climatic changes. The more physically complex coastal landscape that resulted apparently included both new and ancient wetland habitats and associated assemblages. However, the nature of the plant assemblages (= paleoecologic and taphonomic controls) is not yet well elucidated. Testing the hypothesis delineated in this paper and expanding our conclusions will require (1) an evaluation of taphonomic controls in plant-bearing beds, (2) a better understanding of the relation between the plant-bearing strata and their equivalent palynologic zones, (3) correlation between palynologic and paleobotanic data and the sequence-stratigraphic framework already established in other areas, and (4) improvement of the chronostratigraphic chart of the Paraná Basin through the discovery of layers suitable for radiometric dating. ACKNOWLEDGMENTS This research was partly supported by grants from CNPq (RI PQ309322/2007-3, PAS PQ305265/2007-5 and MH PQ300866/2008-9), and FAPERGS (RI 02/1755-2 and PROAPP04/1066.0). Many thanks to two anonymous reviewers who improved the manuscript with many suggestions and helpful comments, and to the editors of this special paper, Luis Buatois and Oscar R. López-Gamundí, for kindly inviting the participation of IGCP Project 471 and for their considerable editorial effort. This is a contribution of the Research Center of Gondwana (CIGO) to the CNPq (process 483463/2007-8) and IGCP Project 471. REFERENCES CITED Anderson, J.M., and Anderson, H.M., 1985, Paleoflora of Southern Africa: Prodromus of South African megafloras, Devonian to Lower Cretaceous: Rotterdam, A. A. Balkema, 423 p. Alves, L.S.R., 1999, Anéis de crescimento em lenhos fósseis como indicadores paleoclimáticos no Neopermiano da Bacia do Paraná (Formação Irati e Formação Serra Alta) [Ph.D. thesis]: Porto Alegre, Instituto de Geociências da Universidade Federal do Rio Grande do Sul. Alves, R.G., and Ade, M.V.B., 1996, Sequence stratigraphy and organic petrography applied to the study of Candiota Coalfield, RS, South Brazil: International Journal of Coal Geology, v. 30, p. 231–248, doi: 10.1016/0166 -5162(95)00041-0. Archangelsky, S., González, C.R., Cúneo, N.R., Sabattini, N., Césari, S.N., Aceñolaza, F.G., Garcia, G.B., Buatois, L.A., Ottone, E., Mazzoni, A.F., Hünicken, M.A., and Gutiérrez, P.R., eds., 1996, El Sistema Permico en la Republica Argentina y en la Republica Oriental del Uruguay: Córdoba, Academia Nacional de Ciencias, 417 p., lam. I–V. Bangert, B., Stollhofen, H., Lorenz, V., and Armstrong, R.L., 1999, The geochronology and significance of ash-fall tuffs in glacigenic, CarboniferousPermian Dwyka Group of Namibia and South Africa: Journal of African Earth Sciences, v. 29, p. 33–49, doi: 10.1016/S0899-5362(99)00078-0. Bortoluzzi, C.A., Piccoli, A.E.M., Corrêa da Silva, Z.C., Cazzulo-Klepzig, M., Dias-Fabrício, M.E., Silva Fo, B.C., Guerra-Sommer, M., Marques-Toigo,

M., Bossi, G.E., and Andreis, R.S., 1980, Estudo geológico da bacia carbonífera de Gravataí-Morungava, RS, in Anais, Congresso Brasileiro de Geologia, 31st, Balneário de Camboriú, Volume 1: Santa Catarina, Brazil, Sociedade Brasileira de Geologia, p. 266–282. Césari, S.N., 2007, Palynological biozones and radiometric data at the Carboniferous-Permian boundary in western Gondwana: Gondwana Research, v. 11, no. 4, p. 529–536, doi: 10.1016/j.gr.2006.07.002. Césari, S.N., and Gutiérrez, P.R., 2000, Palynostratigraphy of Upper Paleozoic sequences in Central-Western Argentina: Palynology, v. 24, p. 113–146, doi: 10.2113/0240113. Chaloner, W., Leistikow, K., and Hill, A., 1979, Brasilodendron gen. nov. and B. pedroanum (Carr.) comb. nov., Permian lycopod from Brazil: Review of Palaeobotany and Palynology, v. 28, no. 2, p. 117–136. Cleal, C.J., 1999, Plant macrofossil biostratigraphy, in Jones, T.P., and Rowe, N.P., eds., Fossil Plants and Spores: Modern Techniques: Geological Society of London, p. 220–224. Daemon, R.F., and Quadros, L.P., 1970, Bioestratigrafia do Neopaleozóico da Bacia do Paraná, in Anais, Congresso Brasileiro de Geologia, 24th, Brasília: Brasília, Sociedade Brasileira de Geologia, p. 359–412. de Matos, S.L.F., Yamamoto, J.K., Riccomini, C., Hachiro, J., and Tassinari, C.C.G., 2001, Absolute dating of Permian ash-fall in the Rio Bonito Formation, Paraná Basin: Gondwana Research, v. 4, p. 421–426, doi: 10.1016/S1342-937X(05)70341-5. Dickins, J.M., 1984, Late Palaeozoic glaciation: BMR Journal of Australian Geology and Geophysics, v. 9, p. 163–169. Gradstein, F.M., et al. (plus 38 authors), 2004, A geologic time scale 2004: Geological Survey of Canada Miscellaneous Report 86, 1 chart. Guerra-Sommer, M., and Cazzulo-Klepzig, M., 1993, Biostratigraphy of the Southern Brazilian Neopaleozoic Gondwana Sequence: A preliminary paleobotanical approach, in Compte Rendu, International Congrès de La Stratigraphie et Géologie du Carbonifère et Permien, 12th, Buenos Aires, 1991, v. 2: Ciudad de Buenos Aires, Argentina, Consejo Nacional de Investigaciones Científicas y Técnicas, p. 61–72. Guerra-Sommer, M., Cazzulo-Klepzig, M., Menegat, R., Formoso, M.L.L., Basei, M.A.S., Barboza, E.G., and Simas, M.W., 2008a, Geochronological data from Faxinal Coal succession in Southern Paraná Basin: A preliminary approach combining radiometric U/Pb age and palynostratigraphy: Journal of South American Earth Sciences, v. 25, p. 246–256, doi: 10.1016/j.jsames.2007.06.007. Guerra-Sommer, M., Cazzulo-Klepzig, M., Santos, J.O.S., Hartmann, L.A., Ketzer, J.M.M., and Formoso, M.L.L., 2008b, Radiometric age determination of tonsteins and stratigraphic constraints for the Lower Permian coal succession in Southern Paraná Basin, Brazil: International Journal of Coal Geology, v. 74, p. 13–27, doi: 10.1016/j.coal.2007.09.005. Holz, M., 1997, Early Permian sequence stratigraphy and paleophysiography of the Paraná Basin in northeastern Rio Grande do Sul State, Brazil: Anais da Academia Brasileira de Ciências, v. 69, no. 4, p. 521–543. Holz, M., 1999, Early Permian sequence stratigraphy and the palaeophysiographic evolution of the Paraná Basin in southernmost Brazil: Journal of African Earth Sciences, v. 29, no. 1, p. 51–61, doi: 10.1016/S0899-5362 (99)00079-2. Holz, M., 2003, Sequence stratigraphy of a lagoonal estuarine system: An example from the Lower Permian Rio Bonito Formation, Paraná Basin, Brazil: Sedimentary Geology, v. 162, no. 3–4, p. 305–331, doi: 10.1016/ S0037-0738(03)00156-8. Holz, M., Vieira, P.E., and Kalkreuth, W., 2000, The Early Permian coal-bearing succession of the Paraná Basin in southernmost Brazil: Depositional model and sequence stratigraphy: Revista Brasileira de Geociências, v. 30, no. 3, p. 420–422. Holz, M., Küchle, J., Philipp, R.P., Bischoff, A.P., and Arima, N., 2006, Hierarchy of tectonic control on stratigraphic signatures: Base-level changes during the Early Permian in the Paraná Basin, southernmost Brazil: Journal of South American Earth Sciences, v. 22, p. 185–204, doi: 10.1016/ j.jsames.2006.09.007. Iannuzzi, R., 2007, Biostratigraphic versus geochronologic frameworks in the Early Permian from Paraná Basin: Looking forward a possible consensus, in Iannuzzi, R., and Boardman, D.R., eds., Extended Abstracts, First Workshop on Problems in Western Gondwana Geology: South America– Africa Correlations: Du Toit Revisited, Gramado, Brazil: Porto Alegre, Centro de Investigações do Gondwana, p. 72–77. Iannuzzi, R., and Souza, P.A., 2005, Floral succession in the Lower Permian deposits of the Brazilian Paraná Basin: An up-to-date overview, in

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The Geological Society of America Special Paper 468 2010

“Levipustula Fauna” in central-western Argentina and its relationships with the Carboniferous glacial event in the southwestern Gondwanan margin Gabriela A. Cisterna* CONICET, Fundación Miguel Lillo, Instituto de Paleontología, Area Geología. Miguel Lillo 251, 4000 Tucumán, Argentina Andrea F. Sterren* CONICET, CIPAL, Centro de Investigaciones Paleobiológicas, Facultad de Ciencias Exactas, Físicas y Naturales, Universidad Nacional de Córdoba. Av. Vélez Sarsfield 299, X5000JJC Córdoba, Argentina

ABSTRACT The Levipustula Fauna (included in the Levipustula levis Zone) is a relatively diversified fossil assemblage composed of brachiopods, bivalves, bryozoans, gastropods, and crinoids. This fauna usually is associated with glaciomarine sequences related to the Carboniferous glacial event that affected the southwestern Gondwanan margin. The Levipustula Fauna has been identified in different units (e.g., Hoyada Verde, La Capilla, Leoncito, and Yalguaraz Formations) exposed in the Calingasta-Uspallata Basin. The Hoyada Verde Formation, herein proposed as a key section, contains the most complete record of the Levipustula Fauna. A detailed compositional, taphonomic, and paleoecological study of this section allows us to propose two associations within the so-called Levipustula Zone: the “Intraglacial Levipustula Fauna,” present in the diamictite-dominated lower part, and the “Postglacial Levipustula Fauna,” dominant in the upper part of section. The fossils of the “Intraglacial Levipustula Fauna” are scarce and poorly diversified. These two features suggest environmentally stressed conditions, probably related to low temperatures in areas close to glaciers. In comparison, the “Postglacial Levipustula Fauna,” relatively more abundant and diverse, exhibits compositional variations that could be explained by paleoenvironmental changes associated with fluctuations in substratum and food supply, such as those identified in modern ecosystems. The identification of the “Intraglacial Levipustula Fauna” and the “Postglacial Levipustula Fauna” may constitute a new tool for understanding the particular relationship between faunal assemblages and climatic variations linked to the Gondwanan glaciation in the Calingasta-Uspallata Basin. Also, the new “Intraglacial Levipustula Fauna” identified in the Hoyada Verde Formation would have biostratigraphical and paleogeographical implications in intra- and interbasinal correlations.

*[email protected]; [email protected] Cisterna, G.A., and Sterren, A.F., 2010, “Levipustula Fauna” in central-western Argentina and its relationships with the Carboniferous glacial event in the southwestern Gondwanan margin, in López-Gamundí, O.R., and Buatois, L.A., eds., Late Paleozoic Glacial Events and Postglacial Transgressions in Gondwana: Geological Society of America Special Paper 468, p. 133–147, doi: 10.1130/2010.2468(06). For permission to copy, contact [email protected]. ©2010 The Geological Society of America. All rights reserved.

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INTRODUCTION The marine invertebrates of the Levipustula Fauna constitute a relatively diversified fossil assemblage composed of brachiopods, bivalves, bryozoans, gastropods, and crinoids. This fauna can be considered the most conspicuous element in different localities in the Calingasta-Uspallata Basin (Fig. 1A), where it usually is associated with glaciomarine deposits. Previous studies (Taboada and Cisterna, 1996; Cisterna, 1999; Cisterna and Ster-

ren, 2004, 2008; Sterren, 2003, 2005) in the classical localities with Levipustula Fauna (Hoyada Verde, Leoncito, and La Capilla Formations) have recognized significant taxonomic, paleoecologic, and taphonomic variations. These variations, as well as the stratigraphic position of this fauna along the glacial sequence, suggest an important paleoenvironmental control. This paper provides a review of the paleogeographic/ paleoclimatic and biostratigraphical framework of the Levipustula Fauna in the Calingasta-Uspallata Basin. In particular, we

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Figure 1. Location maps showing the paleogeography and geography of the Calingasta-Uspallata Basin with the Levipustula Fauna occurrences.

“Levipustula Fauna” in central-western Argentina analyze the relationship of this fauna with the Carboniferous glacial event in the Hoyada Verde Formation, herein proposed as a key section. Because brachiopods and bivalves are the most abundant and diversified invertebrates in the Levipustula Fauna, the present study will be focused on these two groups. PALEOGEOGRAPHIC AND PALEOCLIMATIC SETTING During the late Paleozoic all the Gondwanan subcontinents were affected by an extensive glaciation (Hambrey and Harland, 1981). Paleomagnetic and paleoclimatic data suggest that the pole moved across South America, Southern Africa, Antarctica, and Australia throughout the Carboniferous-Permian transition as a result of the apparent path of polar wander (Crowell, 1983; Caputo and Crowell, 1985; Scotese and Barret, 1990; Scotese and McKerrow, 1990). Three glacial episodes have been suggested within this late Paleozoic paleoclimatic mega-event (López-Gamundí, 1997). In western Gondwana, the earliest glacial episode is recorded in sediments of latest Devonian–earliest Carboniferous exposed in the Solimões and Amazonas Basins of Brazil and the Titicaca Lake region of western Bolivia. The second glacial episode, identified in the Late Carboniferous sediments (Namurian-Westphalian) along the southwestern South American basins, is closely associated with the Levipustula Fauna herein studied. The latest episode has been identified in latest Carboniferous to earliest Permian sediments in basins of eastern South America (Paraná, Sauce Grande, and Malvinas Basins) and South Africa. The late Carboniferous glacial episode is best documented along western Argentina, particularly in the Calingasta-Uspallata Basin, a back-arc basin formed behind an active magmatic arc (Ramos et al., 1986; Ramos and Palma, 1996). A combination of high latitude and altitude allowed the formation of ice centers along the margins of this basin (López-Gamundí, 1997). In this sense, a local north-trending high called the Proto-Precordillera (Amos and Rolleri, 1965), located to the east of the CalingastaUspallata Basin (Fig. 1A), might have exerted an important topographic control during this glacial stage (López-Gamundí and Rosello, 1995). Striated and boulder pavements within the Carboniferous glaciomarine sequences (e.g., Hoyada Verde and Leoncito Formations) have been recognized. Strong northward components in the paleo-ice flow indicators, as well as paleocurrents from the overlying shallow marine sandstones, suggest a common regional paleoslope for both glacial and early postglacial times (López-Gamundí and Martínez, 2000). The Carboniferous glacial sequences of the CalingastaUspallata Basin are characterized by glaciomarine diamictites that grade upward to postglacial open-marine fine clastics. The presence of the marine fine clastics resting on diamictites has been interpreted as the sedimentary response to a glacio-eustatic sea-level rise, expressed as a transgression in stratigraphic terms, that occurred during the glacial retreat subsequent to a widespread glaciation (López-Gamundí, 1989, 1990, this volume).

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The Levipustula Fauna is usually considered to be associated with the postglacial event (López-Gamundí, 1989, 1990; LópezGamundí and Espejo, 1993; López-Gamundí and Martínez, 2000). However, the position of this fauna in the glacial sequence of the Hoyada Verde Formation, as well as its taxonomic and paleoecological/taphonomic variations, suggests the presence of a recently identified “Intraglacial Levipustula Fauna” (Cisterna and Sterren, 2008), besides the typical Levipustula Fauna associated with the postglacial transgression. GEOGRAPHIC AND STRATIGRAPHIC DISTRIBUTION OF THE LEVIPUSTULA FAUNA The Levipustula Fauna has been identified in different diamictite-bearing sections of the Calingasta-Uspallata Basin (Fig. 1B). Detailed studies on the brachiopods and bivalves of this fauna suggest significant taxonomic, taphonomic and paleoecological variations (Cisterna and Sterren, 2003, 2004, 2008; Sterren, 2005). The material studied in this contribution was collected from the classical localities of the Levipustula Fauna of the Hoyada Verde, La Capilla and Leoncito Formations (Table 1). We provide herein a review of the geographic, stratigraphic, and paleontological aspects of these units. Hoyada Verde Formation. The Hoyada Verde Formation (Mésigos, 1953) is exposed in the eroded core of a broad northsouth anticline (the Hoyada Verde anticline) located 3 km east of Barreal village (Fig. 2). The basal contact of the Hoyada Verde Formation is unknown and the top is overlain by the Tres Saltos Formation with an angular unconformity (Fig. 3). The Hoyada Verde section is characterized by a glaciogenic sequence (diamictite and pebbly [dropstone] shale facies) that grades upward into postglacial, dropstone-free shales that contain the typical Levipustula Fauna. An intertill boulder pavement, considered of subglacial origin (López Gamundí and Martínez, 2000), is present in the upper part of the Hoyada Verde Formation and is associated with the diamictic facies. To the uppermost part of the section (Fig. 3), the sequence is characterized by mudstones with interbedded fine sandstones arranged in coarsening and thickening-upward sequences. This interval has been interpreted as the transition from the offshore to the lower shoreface (Buatois and Limarino, 2003), and a succession of fossils traces composed of Psammichnites implexus (Rindsberg) and Psammichnites plummeri (Fenton and Fenton) has been recognized by Mángano et al. (2003). In the top of the Hoyada Verde Formation the uppermost fossiliferous horizon identified is located in a coquina with gastropods such as Peruvispira reedi Sabattini and Mourlonia sp. (Sabattini, 1980). The marine invertebrate assemblage associated with the postglacial shales is composed of brachiopods, bivalves, bryozoans (the fenestellids Fenestella sanjuanesis Sabattini, F. barrealensis Sabattini, F. altispinosa Sabattini, and Polypora neerkolensis Crockfor, described by Sabattini [1972]) and scarce gastropods and crinoids. The brachiopods assemblage is dominated by the Levipustula-Costuloplica-Kitakamithyris association with the

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Cisterna and Sterren TABLE 1. TAXONOMIC COMPOSITION OF THE LEVIPUSTULA FAUNA IN THE CALINGASTA-USPALLATA BASIN

Taxon

HOYADA VERDE FM. Lower part Upper part

Levipustula levis Maxwell Costuloplica leoncitensis (Harrington) Kitakamithyris booralensis (Campbell) Kitakamithyris sp. Beecheria sp. Spiriferellina octoplicata (Sowerby) Spiriferellina sp. Septosyringothyris keideli (Harrington) Orthotetoidea indet. Nuculopsis sp. Phestia sp. aff. P. bellistriata (Stevens) Phestia sp. Palaeolima retifera (Shumard) Streblochondria stappenbecki Reed Streblochondria sanjuanensis Sterren Aviculopecten barrealensis Reed Schizodus sp. Oriocrassatella ? sp. Pleurophorella ? sp. Oriocrassatella andina González Myofossa calingastensis González Leptodesma (Leiopteria) sp. Cypricardinia sp. Promytilus sp. Pyramus ? sp. Naiadites sp. Barrealispira mesigosi Taboada and Sabattini Murlonia striata (Sowerby) Ptychomphalina sabattinii Taboada Ptychomphalina turgentis Taboada Ptychomphalina cf. kuttungensis Taboada Peruvispira reedi Sabattini Straparollus (Euomphalus) sp. Leptoptygma sp. Fenestella sanjuanensis Sabattini Fenestella barrealensis Sabattini Fenestella altispinosa Sabattini Fenestella sp. Polypora neerkolensis Crockford Cladochonus harringtoni Sabattini Sphenotallus stubblefieldi Schmidt and Teichmüller Bryozoans indet. Crinoids indet. Ostracods indet.

C C ------C ------C --C C --------------------------D D ------------------------D C C C

D D --D D D --------D --D D D --------------------------------D ----D D D --D D ----C ---

LA CAPILLA FM. North (LJ) South (LC) D D D D D M D --------D ------D D D D --------------------------------------------M M ---

--M D D M M M M M --------------M ----D D D D D D ------D D D D M M --------------M -----

LEONCITO FM. M D ----D D --D M M --M --------M --M ------------D M --------------------M ------M M M

C - Cited in this work D - Described in previous work M - Mentioned in previous work

species Levipustula levis Maxwell, Costuloplica leoncitensis (Harrington), and Kitakamithyris sp., usually accompanied by Beecheria sp. and “Spiriferellina” octoplicata (Sowerby). The bivalve fauna studied by Sterren (2003) is composed of Phestia sp. aff. P. bellistriata (Stevens), Palaeolima retifera (Shumard), Streblochondria sanjuanensis Sterren, and S. stappenbecki Reed. Although this is the main faunal assemblage in the Hoyada Verde

Formation, the gastropods Neilsonia? sp., Neoplatyteichum barrealense (Reed), Barrealispira mesigosi Taboada and Sabattini, Ptychomphalina striata (Sowerby), the annelid Sphenotallus stubblefieldi Schmidt and Teichmüller, and the brachiopods Levipustula levis Maxwell and Kitakamithyris sp. have been identified in the lower part of the section (Sabattini, 1980; Taboada and Sabattini, 1987; Taboada, 1997). Taboada (1997) has also

“Levipustula Fauna” in central-western Argentina

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69°25´ W

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Figure 2. Generalized geological map showing the distribution of the Hoyada Verde Formation outcrops. Modified from Mésigos (1953).

suggested the presence of brachiopods, probably the species Levipustula levis Maxwell, associated with the diamictite-rich glacial section. The Hoyada Verde Formation has been considered chronologically equivalent to the El Paso Formation (Amos and Rolleri, 1965; Amos and López-Gamundi, 1981; González, 1990), a diamictite-rich section exposed in the southernmost part of the Barreal hill (Fig. 2). However, there no physical continuity between the two units, and the brachiopods identified in the El Paso Formation (Micraphelia indianae Simanauskas and Cisterna, Tuberculatella peregrina (Reed), Aseptella aff. A. patriciae Simanauskas, and Rhipidomella? sp.) would exhibit latest Carboniferous biostratigraphical affinities (Simanauskas and Cisterna, 2001). La Capilla Formation. Outcrops of the La Capilla Formation (Amos et al., 1963) have been recognized in two sectors in the Calingasta area (Fig. 1B). One of these exposures

is located 1400 m north La Capilla village, along the road that connects Calingasta and the Castaño Viejo mines, ~300 m west of Las Cambachas. This 39-m-thick section is essentially composed of sandstones, gray-green mudstones, coquinoid lenticular beds, and scarce fine conglomerates (Amos et al., 1963). The Levipustula Fauna occurs in the upper part of the section and is taxonomically close to that present in the Hoyada Verde Formation. However, the taxonomic differences among bivalves (Phestia sp., Aviculopecten barrealensis Reed, Schizodus sp., Oriocrassatella? sp., and Pleurophorella? sp.) are more important than those of the brachiopods (Cisterna and Sterren, 2003, 2004; Sterren, 2005). Other outcrops of the La Capilla Formation are located ~5 km north of La Capilla village, in a locality known as Las Juntas. Sessarego and Amos (1987) recognized two members: a lower member, where the Levipustula Fauna occurs, has been interpreted as deposited in a proximal glaciomarine setting

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Tres Saltos Fm.

PG

Shales Laminated shales with dropstones Fine-grained sandstones

Figure 3. Stratigraphic section of the Hoyada Verde Formation (modified from López-Gamundí, 1983) and vertical distribution of the Intraglacial (IG) and Postglacial Levipustula Fauna (PG).

Stratified pebbly mudstone Massive bouldery to pebbly sandy mudstone Striated boulder pavement IG

Levipustula fauna

20 m

Gastropods Ichnofossils

?

with deltaic influence (Sessarego and Amos, 1987), and an upper member, consisting of deltaic greenish brown sandstones and mudstones (Vallecillo and Bercowski, 1998). The Levipustula Fauna in this section is not well known. The brachiopod assemblage appears to be dominated by the genus Kitakamithyris, with Kitakamithyris booralensis (Campbell) and Kitakamithyris sp., accompanied by Septosyringothyris aff. S. keideli (Harrington) and Orthotetoidea indet. González and Taboada (1987) and González (2002) have also suggested the presence of the brachiopods Spiriferellina octoplicata (Sowerby), Costuloplica leoncitencis (Harrington), Septosyringothyris keideli (Harrington), Kitakamithyris septata? (Chronic), Lingula sp., and Chonetacea indet.; the bivalves Oriocrassatella andina González, Myofossa calingastensis González, Leptodesma (Leiopteria) sp., Schizodus sp., Promytilus sp., Pyramus? sp., and Cypricardinia? sp.; the gastropods Peruvispira reedi Sabattini, Peruvispira cf. kuttungensis Campbell, Murlonia ssp., Straparollus (Euomphalus) sp., Leptoptygma sp.; and the bryozoans Fenestella sp.

Leoncito Formation. Outcrops of the Leoncito Formation (Baldis, 1964) are located ~22 km southeast of Barreal village on the western flank of the Precordillera (Fig. 1B), along the southern margin of the Las Cabeceras river. This section is dominated by sandy facies and the diamictite beds appear to the top of the section, where a striated pavement has been also identified. This pavement is shaped on bioturbated fine-grained sandstones with plant fragments and the surface is covered by a massive diamictite (López-Gamundí and Martínez, 2000). In the Leoncito Formation section the Levipustula Fauna occurs in sandstone and mudstone horizons, located below the glacial diamictic beds in an interval ~11 m thick. The fauna associated with the sandstones appears concentrated in distinct lenses and is composed of the Septosyringothyris keideli–Costuloplica leoncitensis brachiopod assemblage, accompanied of “Spiriferellina” octoplicata Sowerby and very scarce Beecheria sp. and Levipustula levis Maxwell; bivalves (Phestia sp., Schizodus sp. and Pleurophorella? sp.), bryozoans (Fenestella? sp.); and gastropods (Barrealispira?

“Levipustula Fauna” in central-western Argentina sp.). The fauna from the mudstone-dominated interval is quite different and composed of brachiopods (Levipustula levis and Beecheria sp.), bivalves (Nuculopsis? and bivalvia indet.), ostracods (probably some Aurykirkbya), gastropods, crinoids, and bryozoans (Cisterna and Sterren, 2008). Although we have not conducted fieldwork in the Yalguaraz Formation outcrops, we are tempted to state that this unit contains some diagnostic elements of the Levipustula Fauna. The Yalguaraz Formation (Amos and Rolleri, 1965) is exposed on the west flank of the Cordillera del Tigre, close to the boundary between San Juan and Mendoza provinces (Fig. 1B). In the type section (Arroyo del Tigre Creek), this unit is characterized by a predominantly diamictic sequence with mudstones and sandstones increasing toward the upper part. An intertill striated pavement on the top of this section has been also suggested (Taboada, 1997). The marine fossil assemblage, present in mudstones and sandstones of the middle part of the section, is composed of brachiopods, gastropods, bryozoans and bivalves (Taboada and Carrizo, 1992; Taboada, 1997). From this assemblage, Taboada and Cisterna (1996) have described the brachiopods Kitakamithyris immensa (Campbell) and Torynifer tigrensis Taboada and Cisterna, which can be considered conspicuous elements of the Levipustula Fauna. AGE AND FAUNAL AFFINITIES WITH OTHER BASINS The Levipustula Fauna is known from eastern Australia where it is also characterized by a low-diversity, cold-water assemblage dominated by brachiopods referred to the Levipustula levis Zone (Campbell and McKellar, 1969; Jones et al., 1973). It has not been associated with a specific reference section within eastern Australian basins but Roberts et al. (1976) have provided a summary of its composition, stratigraphic occurrence, and relationships with other zones. The age of this zone was widely discussed by different authors (Campbell, 1961; McKellar, 1965; Lindsay, 1969; Jones et al., 1973; Roberts, 1976; Roberts et al., 1976, 1993) and finally assigned, on the basis of faunal considerations, to the Namurian–Westphalian interval (Roberts et al., 1976). However, sensitive high-resolution ion microprobe (SHRIMP) zircon dating of volcanic horizons in the glaciogenic sediments of the Seaham Formation (Southern New England Orogen) appears to confine the age of this zone to early Namurian (Roberts et al., 1995). In Argentina the Levipustula levis Zone was defined by Amos and Rolleri (1965) in the Calingasta-Uspallata Basin. The La Capilla Formation has been proposed as the reference section of the Levipustula levis Zone, but a detailed study in the Hoyada Verde Formation suggests that the latter unit contains the most complete record of the Levipustula Fauna in the basin. The age of the Levipustula levis Zone in Argentina was considered Namurian-Westphalian owing to its Australian affinity (González, 1981; Archangelsky et al., 1987; Archangelsky et al., 1996; Taboada, 1997). The lower limit of age for the Levipustula

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levis Zone in Precordillera was discussed by Taboada (1997), who recognized a mudstone-rich interval in the lowermost Hoyada Verde Formation with Barrealispira mesigosi Taboada and Sabattini, Ptychomphalina striata (Sowerby) and Sphenotallus stubblefieldi Schmidt and Teichmüller, associated with Levipustula levis Maxwell. Such an assemblage would indicate an early Namurian age. A recent biostratigraphic review of the Hoyada Verde Formation (Sterren and Cisterna, 2006) indicates that the fauna suggested by Taboada (1997) occurs within the diamictic section. There are no diagnostic elements to suggest the precise upper limit of this fauna but, for the moment, its biostratigraphic relationships would indicate a Westphalian age (González, 1981, 1990, 1993; Archangelsky et al., 1987). Brachiopods of the Levipustula levis Zone in the Precordillera of Western Argentina appear to be the main tool for biostratigraphic correlation, and their Australian faunal affinities would be essentially based on the common species Levipustula levis Maxwell, Kitakamithyris booralensis (Campbell), and Kitakamithyris immensa (Campbell) (Taboada and Cisterna, 1996). Although there are previous systematic studies (Reed, 1927; Keidel and Harrington, 1938; Amos et al., 1963; Lech, 1989; Taboada and Cisterna, 1996), a taxonomic review of new collections of Levipustula Fauna brachiopods from Precordillera started by one the authors (G.A.C.) can shed further light on some of the current correlation problems. In this sense, the first problem related to the systematics of this fauna’s brachiopods is the original definition of Levipustula levis in Australia (Maxwell, 1951). This species was described from different units in eastern Australia, such as the Booral Formation in New South Wales (Campbell, 1961) and the Poperima Formation and Branch Creek Formation in Queensland (Maxwell, 1964). However, specimens from different localities assigned to Levipustula levis appear to have distinct species-diagnostic features. Other South American basins where the Levipustula Fauna have also been identified are the Tepuel-Genoa Basin in southwestern Argentina and the Tarija Basin in Bolivia. In the Tepuel-Genoa basin, the Levipustula levis Zone was previously recognized in the Pampa de Tepuel and Las Salinas Formations (Amos et al., 1973). However, the brachiopods from the Tepuel-Genoa basin originally assigned to the Levipustula levis Maxwell by Amos (1960) were included in the synonymy of Lanipustula patagoniensis Simanauskas (1996) and Verchojania inacayali Taboada (2008). The genus Lanipustula proposed by Klets (1983) is very close to Levipustula Maxwell but it has been differentiated by the disposition of the cardinal ridges and the shape of the anterior adductor muscle scars (Simanauskas, 1996). Recent studies of the genera Levipustula and Lanipustula in Argentina suggest that the diagnostic features proposed for distinguishing these genera by Klets (1983) can be considered of intraspecific hierarchy (Taboada, 2006). Instead, other main differences, such as the abundance of external dorsal spines and variable rugae and/or lamination on both valves developed in Lanipustula, clearly separate it from Levipustula (A.C. Taboada, 2010, personal commun.).

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A multivariate analysis of the Levipustula levis Zone in the Tepuel-Genoa Basin was conducted by Simanauskas and Sabattini (1997), who subdivided the zone into the late Carboniferous Lanipustula Zone, the Early Permian Pyramus faunule (Asselian), and the Tuberculatella Zone (Sakmarian). The Lanipustula Zone was referred to the lower part (“Fenestella and Productus” horizon) of the Pampa de Tepuel Formation (Freytes, 1971) and to the lower member of the Las Salinas Formation (González, 1972); Simanauskas and Sabattini (1997) suggested a Namurian– Stephanian age for this zone. In the Tarija Basin the Levipustula Fauna is not well known but an invertebrate marine assemblage composed of Levipustula levis Maxwell, Cypricardinia? boliviana Rocha-Campos, Carvalho and Amos, Limipecten cf. L. burnettensis Maxwell, Stuchburia sp., Myonia sp., and Mourlonia balapucense RochaCampos, Carvalho and Amos was described from the upper part of the Taiguatí Formation (Bolivia) by Rocha-Campos et al. (1977). The bivalves Naiadites cf. N. modiolaris (Sowerby) and Wilkingia cf. W. elliptica (Phillips) have been identified from the Taiguatí Formation as well (Trujillo Ikeda, 1989). LEVIPUSTULA FAUNA IN THE HOYADA VERDE FORMATION The typical Levipustula Fauna can be easily identified associated with mudstone facies, located above of the glacial diamictic sequence, in the upper part of the Hoyada Verde Formation. However, a detailed study along this section indicates that a scattered and very poorly diversified faunal assemblage can be recognized in the lower part of the section, interbedded with diamictic horizons. These faunas are herein proposed as the “Postglacial Levipustula Fauna” (Fig. 3, PG) and “Intraglacial Levipustula Fauna” (Fig. 3, IG), respectively, and their compositional, taphonomic, and paleoecological features are discussed. The “Intraglacial Levipustula Fauna” The outcrops that contain the “Intraglacial Levipustula Fauna” can be recognized in the exposed core of the Hoyada Verde anticline. Therefore, these mudstones with fauna have been usually referred to the lowermost part of the section (Mésigos, 1953; Taboada, 1997). Recent field work has allowed us to recognize the correct location of this fossiliferous interval and its stratigraphic relationships (Fig. 3). The new fossil assemblage appears in the lower part of the Hoyada Verde Formation, interbedded with diamictic horizons, in a 20-m-thick interval made up of grayish, laminated mudstones characterized by the presence of glendonites and dropstones. The “Intraglacial Levipustula Fauna” (Fig. 4) is characterized by a monotonous assemblage of marine invertebrates and stems of “Dadoxilon.” The fauna is dominated by brachiopods, bivalves, and annelids, accompanied by gastropods, ostracods, and frag-

mentary bryozoans. The brachiopods recognized are Levipustula levis? Maxwell and Levipustulini indet., apparently the dominant species, and very scarce specimens of “Spiriferellina” octoplicata and Spiriferidae indet. The bivalves distinguished in this fauna are Nuculopsis sp., Phestia sp. aff. P. bellistriata (Stevens) and Palaeolima retifera (Shumard). A preliminary study of the fauna associated with the brachiopods and bivalves has allowed us to recognize the ostracods Kirkbyidae indet. (probably the genus Aurykirkbya Sohn), the annelids Sphenotallus stubblefieldi Schmidt and Teichmüller, the gastropods Barrealispira sp., and the bryozoans Fenestella? sp. The intraglacial fauna is scarce, poorly diversified, and very scattered within thick mudstone packages. The bioclasts show a random distribution, degree of fragmentation is low, and delicate details of the fine sculpture are preserved. In addition, individual shells are generally small, exhibiting a relatively wide range of valve sizes, from 0.3 to 2 cm. Shells of both brachiopods and bivalves are mostly disarticulated; however, specimens belonging to the bivalve family Nuculanidae (Nuculopsis sp., Phestia sp. aff. P. bellistriata (Stevens)), are exceptionally found with articulated valves. The attributes described for these fossils suggest a fauna buried in situ. Features such as its low diversity and abundance indicate environmentally stressed conditions probably related to glacial environment. The presence of articulated bivalves might be related to minor reworking and transport, and conditions of sudden burial (Kidwell and Bosence, 1991; Aigner, 1985; Peterson, 1985) when the ligament was still active (Fürsich and Heinberg, 1983). In addition, low temperatures might have delayed the decomposition of the soft tissues that unite the valves (Kidwell and Baumiller, 1990). The bivalves that characterize this interval, Nuculopsis and Phestia, are commonly found associated with stress conditions (Sterren, 2000; Simanauskas and Cisterna, 2000; Lebold and Kammer, 2006). The abundance of the eurytopic bivalve Nuculopsis is significant because the deposit-feeding nuculid bivalves are common component of fossil assemblages in oxygen-deficient basins (Kammer et al., 1986). The annelid Sphenotallus stubblefieldi is another conspicuous element in the “Intraglacial Levipustula Fauna.” The annelids are considered stress-tolerant taxa and their presence in glaciomarine sequences indicates a wide range of temperature, salinity, and oxygen tolerance (Collinson et al., 1994; Chakraborty and Bhattacharya, 2005). The presence of glendonite concretions and dropstones in the stratigraphic interval that contains the “Intraglacial Levipustula Fauna” suggests extreme environmental conditions related to the glacial proximity. The glendonite, a carbonate pseudomorph that might indicate freezing water (González, 1980; Swainson and Hammond, 2001; McLachlan et al., 2001) has been also recorded in similar Carboniferous and Permian sequences of Argentine Patagonia (Tepuel-Genoa basin, González, 1980), as well as in South Africa and Australia (McLachlan et al., 2001; Thomas et al., 2007; López-Gamundí, this volume).

B A

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Figure 4. “Intraglacial Levipustula Fauna.” (A–D, G) Levipustula levis? Maxwell. (A, G) internal and external mold of ventral valve, CEGH-UNC 22171 (×3); (B) internal mold of ventral valve, CEGH-UNC 22172 (×2.5); (C) external mold of ventral valve, CEGH-UNC 22173 (×3); (D) internal mold of ventral valve, CEGH-UNC 22175 (×2.5); (E) Levipustululini indet., internal mold of ventral valve, CEGH-UNC 22174 (×2.5); (F) “Spiriferellina” octoplicata (Sowerby) fragmentary ventral valve, CEGH-UNC 22176 (×4); (H) Phestia sp. aff. P. bellistriata (Stevens) outer view of left valve, CEGH-UNC 22160 (×9); (I) Palaeolima retifera (Shumard) outer view of right valve, CEGH-UNC 22161 (×4). (J–L) Nuculopsis sp. (J) internal mold of left valve, CEGH-UNC 22163 (×10); (K) internal mold of articulated valves, CEGH-UNC 22162 (×14); (L) interior of left valve, CEGH-UNC 22164 (×11). Fossils with the prefix CEGH-UNC are housed in the Centro de Investigaciones Paleobiológicas (Universidad Nacional de Córdoba) and those with the prefix IPI are housed in the Fundación Miguel Lillo (Instituto de Paleontología).

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The “Postglacial Levipustula Fauna” The first marine invertebrates of the “Postglacial Levipustula Fauna” have been identified ~20 m above of the uppermost diamictic horizon of the Hoyada Verde Formation (Fig. 3). The fossil assemblages associated with the postglacial shales are composed of bryozoans, brachiopods, and bivalves (Fig. 5), accompanied by less abundant gastropods and crinoids. This fauna is very abundant and highly diversified, and it exhibits compositional variations throughout the fossiliferous interval studied. The fossil concentrations occur either as thin (1–5 cm) shell beds or nests. The bioclasts show poor sorting and a random distribution. In cross section, the shells are mainly concordant to slightly oblique to the bedding plane and display predominantly concave-upward orientations. Low degrees of abrasion and fragmentation have been observed in the shells. Similar proportions of dorsal/ventral valves in brachiopods and left/right valves in bivalves are recognized. Some articulated shells are present and delicate details of fine sculpture, such as spines in Levipustula levis, are also preserved. The taphonomic features described suggest biogenic fossil concentrations, produced by a gradual accumulation of successive benthic colonizations (Sterren, 2002). Cisterna (1999) carried out a detailed paleoecologic analysis that includes the dominant groups (bryozoans, brachiopods and bivalves) of the postglacial fauna. From this paleoecologic study, Simanauskas et al. (2001) recognized three subfaunas (Fig. 6), based primarily on the changes observed in the brachiopods in the fossiliferous interval: the lower Costuloplica leoncitensis subfauna, an intermediate Kitakamithyris sp. subfauna, and the upper, Levipustula levis subfauna. From our most recent fieldwork we were able to confirm the presence of the three brachiopod associations previously identified by Simanauskas et al. (2001) in the postglacial interval (although the bryozoans and the brachiopod Costuloplica leoncitensis appear to be conspicuous elements throughout the interval). The lower association (Fig. 6A) is characterized by the low number of species and low total biovolume. The communities of this association are dominated by bivalves of the genus Streblochondria, accompanied, in decreasing order, by bryozoans and Costuloplica leoncitensis, organisms characterized as epifaunal and suspension feeding. This association would represent a gradual deepening of cold waters and a relatively low nutrient availability, which is evidenced by the low biovolume and the low species richness. The intermediate association (Fig. 6B) is dominated by brachiopods or bryozoans and characterized by the highest species richness and the highest values of biovolume. The bivalves are less important in these communities but those of epibyssate habits (Streblochondria sanjuanensis Sterren, Streblochondria stappenbecki Reed, and Palaeolima retifera (Shumard)) show a relative increase. The brachiopod communities of this association are dominated by Costuloplica leoncitensis, accompanied by Kitakamithyris sp., “Spiriferellina” octoplicata, Beecheria sp., and Levipustula levis as a subordinate element. This brachiopod association could correspond to the maximum flooding with a substra-

tum more stable in which the suspension-feeding and epifaunal organisms appear more diversified (epibyssate, libero-sessile, and pedunculate habits). The relatively more benign climatic conditions would have triggered an abundant food supply. Hence, the communities of this association have the highest faunal species richness and biovolume. The fossiliferous interval that contains the upper brachiopod association is characterized by abundant sandstones linked to a shallowing-upward trend. This association is dominated by Levipustula levis, a small quasi-infaunal productid, and the ubiquitous Costuloplica leoncitensis accompanied by the bivalve Phestia sp. aff. P. bellistriata (Fig. 6C). The values of species richness and biovolume are relatively low and the presence of the opportunistic organisms Levipustula levis and Phestia sp. aff. P. bellistriata could be related to fluctuating environmental conditions (i.e., an unstable substratum and higher sediment rate). The communities of the three associations described above would have been developed in a stable marine environment, such as an open shelf with moderate bottom currents (Cisterna, 1999). Variations in these associations would have been controlled by substrate types and food supply fluctuations during the postglacial transgression. DISCUSSION A complex glacial history with advances and retreats of glaciers might have been the main control on the distribution of the Levipustula Fauna. Although this fauna is a conspicuous element in the Calingasta-Uspallata Basin, the glacial influence is more evident in the faunas present in the Hoyada Verde Formation. The identification of the “Intraglacial Levipustula Fauna” and the “Postglacial Levipustula Fauna” described in this paper constitutes a new element for understanding the particular relationship between the faunal assemblages and the climatic variations due to the Gondwanan glaciation. The features discussed for the Intraglacial Fauna strongly suggest environmentally stressed conditions probably related to low (glacial) temperatures. Although low temperatures seem to be the dominant factor of stress for the faunas discussed herein, other variables such as oxygen and nutrient availability, salinity, substrate type, and water depth would have been affected by the glacial action. The very abundant and highly diversified “Postglacial Levipustula fauna” can be considered to represent the record of climatic amelioration in more stable environmental conditions. This fauna exhibits compositional variations that reflect slight changes in the substrate stability and food supply, related to the moderate bottom currents. Some features of this fauna in the middle part of the postglacial interval, such as high diversity of the filter-feeder and epifaunal organisms, suggest the presence of maximum flooding conditions during the postglacial transgression. The major sandstone component of the uppermost part of the interval is related to a shallowing-upward trend. The low

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Figure 5. Postglacial “Levipustula Fauna.” (A–C) Levipustula levis Maxwell. (A) dorsal valve partially decorticated, IPI 3442 (×2.5); (B) internal mold of dorsal valve, IPI 4501 (×3); (C) ventral valve, IPI 3238 (×2). (D–E) Costuloplica leoncitensis (Harrington); (D) ventral valve, IPI 3221 (×1); (E) dorsal valve partially decorticated, IPI 3223 (×2.5). (F–G) Kitakamithyris sp. (F) Internal mold of ventral valve incomplete, IPI 4502 (×1); (G) internal mold of dorsal valve, IPI 4503 (×1). (H–I) “Spiriferellina” octoplicata (Sowerby); (H) ventral valve incomplete, IPI 3232 (×3.5); (I) internal mold of dorsal valve, IPI 3235 (×2.3). (J–K) Beecheria sp.; (J) internal mold of ventral valve, IPI 4504 (×5); (K) internal mold of dorsal valve, IPI 4505 (×4). (L–M) Streblochondria sanjuanensis Sterren; (L) paratype, composite mold of left valve, CEGH-UNC 19750 (×1.5); (M) holotype, composite mold of right valve, CEGH-UNC 19748 (×1.5). (N–O) Streblochondria stappenbecki (Reed); N, internal mold of left valve, CEGH-UNC 19757 (×2.5); (O) internal mold of right valve, CEGH-UNC 19755 (×1.8). (P–Q) Palaeolima retifera (Shumard); (P) outer view of right valve, CEGH-UNC 22168 (×7); (Q) outer view of right valve, CEGH-UNC 19773 (×2.3). (R–S) Phestia sp. aff. P. bellistriata (Stevens); (R) interior of left valve, CEGH-UNC 22165 (×4); (S) dorsal view of right valve showing dentitions, CEGH-UNC 19717 (×6).

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Figure 6. Hypothetical paleoecological reconstructions of the three brachiopod associations recognized in the postglacial interval. (A) Lower association dominated by bivalves of the genus Streblochondria, accompanied of bryozoans and Costuloplica leoncitensis. (B) Middle association dominated by brachiopod communities, characterized by the highest species richness and the highest values of biovolume. (C) Upper association dominated by Levipustula levis and Costuloplica leoncitensis, accompanied by Phestia sp. aff. P. bellistriata. a—Costuloplica leoncitensis (Harrington), b—Levipustula levis Maxwell, c—Beecheria sp., d—“Spiriferellina” octoplicata (Sowerby), e—Kitakamithyris sp., f— Streblochondria sanjuanensis Sterren, g—Streblochondria stappenbecki Reed, h—Palaeolima retifera (Shumard), i— Phestia sp. aff. P. bellistriata, j—Fenestella sp., k—Crinoidea indet.

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diversity of the fauna in this part of the postglacial interval as well as the presence of the opportunistic organisms may represent a relatively unstable substratum. The distinctive pattern of the faunal distribution in the Hoyada Verde Formation, as well as the persistence of some taxa after the glaciation, have been recognized and studied in modern ecosystems close to glaciers. Faunal variations in the postglacial phases, mainly related to differences in substrate types, water depth, and variations in clastic and organic content sediment rates, have been suggested by Gordillo and Aitken (2001) for the modern ecosystems of the Arctic region. These important changes are reflected in the development of different associations, from the onset of the deglaciation to the postglacial phase. When the glacier retreats, large volumes of sediment discharged by the meltwater flows can suppress the vertical circulation and the nutrient generation in the surface water, limiting the marine primary production and the food supply in ice-proximal benthic habitats. These condi-

tions drastically change in the postglacial phase and the increase of the nutrients accounts for the diversification of the mollusk assemblages. This model can be useful for understanding why the “Postglacial Levipustula Fauna” is more diversified than the “Intraglacial Levipustula Fauna”; it also allows us to understand the other variations of the postglacial fauna in the different associations identified in this contribution. Studies of the behavior of modern bivalve assemblages from circumpolar regions suggest that during a glacial event the ice can negatively affect much of the benthic marine fauna in continental shelves, but some species are able to survive using some form of refugium. Various strategies for the subfossil mollusks have therefore been proposed, such as displacing into deeper waters, moving into subpolar regions, or surviving in nonglaciated pockets in continental shelves (Crame, 1996). This type of behavior in glacial conditions could explain the recurrence of some bivalves along the Hoyada Verde section, such as the presence of Phestia

“Levipustula Fauna” in central-western Argentina sp. aff. P. bellistriata and Palaeolima retifera in the Intra- and Postglacial Faunas. The influence of the glaciation in the late Paleozoic marine biota has been also documented in other basins of western Argentina such as the western Paganzo Basin. Pazos (2000) recognized a glacial opportunistic ichnofaunal assemblage characterized by a low diversity and a high degree of burial, which could suggest stress conditions in a cold-water environment. Buatois et al. (2006, this volume) proposed that high sedimentation rates and fluctuations in the water salinity might have characterized the depositional conditions during the Gondwanan glaciation. These authors suggested in addition that the postglacial assemblages were developed during a transgressive event and the subsequent deltaic progradation, which set new conditions characterized by abundant supply of nutrients and oxygenation of the water column. The conclusions from the present study conducted in the Hoyada Verde Formation can be extended to other areas of the Calingasta-Uspallata Basin. Although the Hoyada Verde section exhibits the best record of the glacial-postglacial transition and associated faunas, a fossil assemblage compositionally and taphonomically equivalent to the “Intraglacial Levipustula Fauna” has been recently recognized in the Leoncito Formation (Cisterna and Sterren, 2008). Ostracods are unusual in the Carboniferous sequences of the basins of western Argentina, and the presence of the genus Aurykirkbya Sohn in the “Intraglacial Levipustula Fauna” could have significant biostratigraphic implications. This ostracod genus has been described from equivalent stratigraphic sequences in the Tepuel-Genoa Basin in Argentine Patagonia. Díaz Saravia and Jones (1999) considered that the ostracod faunas described from this basin in the lower part of the Pampa de Tepuel Formation would be early Namurian (lower part of the Levipustula levis Zone). In this sense the new “Intraglacial Levipustula Fauna” identified in the Hoyada Verde Formation would have biostratigraphic and paleogeographic implications for intraand even interbasinal correlations. ACKNOWLEDGMENTS The authors would like to thank CONICET, Consejo Nacional de Investigaciones Científicas y Técnicas (PIP 6144), and the Agencia Nacional de Promoción Científica y Tecnológica (PICT 20752-PICT 32693) from Argentina. We acknowledge N. Emilio Vaccari for his help in the photographic work. The reviewers, Lucia Angiolini and Marcello Simões, improved an earlier version of the manuscript. REFERENCES CITED Aigner, T., 1985, Storm depositional systems: Dynamic stratigraphy in modern and ancient shallow marine sequences: Lecture Notes in Earth Sciences, v. 3, p. 1–174, doi: 10.1007/bfb0011412. Amos, A.J., and López Gamundi, O.R., 1981, Late Paleozoic diamictites of the Calingasta-Uspallata and Paganzo basins, San Juan and Mendoza provinces, in Hambrey, M., and Harland, W., eds., Earth’s Pre-

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Printed in the USA

The Geological Society of America Special Paper 468 2010

Ichnology of late Paleozoic postglacial transgressive deposits in Gondwana: Reconstructing salinity conditions in coastal ecosystems affected by strong meltwater discharge Luis A. Buatois Department of Geological Sciences, University of Saskatchewan, 114 Science Place, Saskatoon, SK S7N 5E2, Canada Renata G. Netto Unisinos, Programa de Pós-graduação em Geologia, Av. Unisinos 950, 93022-000 São Leopoldo RS, Brazil M. Gabriela Mángano Department of Geological Sciences, University of Saskatchewan, 114 Science Place, Saskatoon, SK S7N 5E2, Canada

ABSTRACT Late Paleozoic ichnofaunas from eight different Gondwanic basins (Paganzo, San Rafael, Tarija, Paraná, Karoo, Falkland, Transantarctic, and Sydney) provide valuable evidence for reconstructing the environmental conditions of postglacial transgressions. The depositional environment of most of these transgressive finegrained deposits historically has been controversial, with interpretations ranging from freshwater lacustrine to brackish-water estuarine, and even normal-salinity, open-marine platforms. Although the various units differ in the degree of marine connection, the common theme in all is the presence of freshwater ichnofaunas in direct association with glacially influenced coasts affected by strong discharges of meltwater. Ichnofaunas are typically dominated by nonspecialized grazing trails (Mermia, Helminthopsis, Helminthoidichnites), simple feeding traces (Treptichnus), arthropod trackways (Diplichnites, Umfolozia), and fish trails (Undichna), representing examples of the Mermia and, to a lesser extent, the Scoyenia ichnofacies. A complex paleogeography of fjords and deep, large coastal lakes is suggested. Freshwater conditions were prevalent during most of the time because these areas were affected by a strong discharge of fresh water due to melting of the ice masses during deglaciation. The simple dichotomy between marine and nonmarine settings is misleading because these peculiar assemblages should first be understood in terms of their paleoecologic significance, and subsequently placed within a larger paleoenvironmental context. Laterally persistent, albeit diachronous, peri-Gondwanan ichnofaunas characterize melting of the late Paleozoic ice caps. Temporal recurrence of these ichnofaunas through the Late Carboniferous–Middle Permian indicates a common response of benthic faunas under similar ecological conditions during deglaciation events.

Buatois, L.A., Netto, R.G., and Mángano, M.G., 2010, Ichnology of late Paleozoic postglacial transgressive deposits in Gondwana: Reconstructing salinity conditions in coastal ecosystems affected by strong meltwater discharge, in López-Gamundí, O.R., and Buatois, L.A., eds., Late Paleozoic Glacial Events and Postglacial Transgressions in Gondwana: Geological Society of America Special Paper 468, p. 149–173, doi: 10.1130/2010.2468(07). For permission to copy, contact [email protected]. ©2010 The Geological Society of America. All rights reserved.

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INTRODUCTION Upper Paleozoic rocks in the Southern Hemisphere contain an extensive record of the Gondwana glaciations and the associated postglacial transgressions (see López-Gamundí, this volume). The third Gondwanan Ice House Age spanned at least 90 m.y. (Caputo and Crowell, 1985), starting in western South America during the Tournaisian (Mississippian) and ending at the base of the Guadalupian (Middle Permian) in Australia. It characterizes the longest continuous glaciation during the Phanerozoic (Eyles, 1993). Upper Paleozoic glacial diamictites and postglacial fine-grained deposits occur in Australia, Southeast Asia, India, Pakistan, Arabian Peninsula, North Africa, Namibia, South Africa, southern South America, and Antarctica (Fig. 1) (e.g., Veevers and Powell, 1987; López Gamundí, 1989, 1997; Eyles and Young, 1994; Visser, 1997; Limarino et al., 2002; Angiolini

et al., 2003; Isbell et al., 2003a, 2003b; Trosdtorf et al., 2005). Deglaciation sequences are diachronous, reflecting the different ages of the Gondwana glaciation. Although Gondwana glacial events and associated deglaciation episodes may have strongly affected benthic communities inhabiting terrestrial, freshwater, and coastal ecosystems, surprisingly little is known about late Paleozoic glacial ecosystems. The global impact of the Gondwana glaciation in marine ecosystems was recently evaluated by Stanley and Powell (2003) and Powell (2005), based on J.J. Sepkoski’s (2002) database and an expanded database of brachiopod occurrences, respectively. These authors noted that rates of origination and extinction for marine organisms dropped to low levels during the glaciation. However, the direct impact of glacial and deglaciation episodes on late Paleozoic Gondwana communities has remained essentially underexplored. In part, lack of studies may be due to the

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Discussed Gondwanan basins with Upper Paleozoic glaciogenic successions (and postglacial freshwater ichnofaunas): 1 - Tarija Basin 2 - Paganzo Basin 3 - San Rafael Basin 4 - Paraná Basin (including Chaco-Paraná and North Uruguayan Basins)

5 - Karoo Basin (including Kalahari Basin) 6 - Falkland Basin 7 - Transantarctic Basin 8 - Sydney Basin

Other Gondwanan glacial basins with Upper Paleozoic glaciogenic successions

Figure 1. Location map of late Paleozoic Gondwanan basins (modified from Isbell et al., 2003a).

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Ichnology of late Paleozoic postglacial transgressive deposits in Gondwana low preservation potential of body fossils in areas adjacent to glaciated margins, such as modern fjords (e.g., Aitken, 1990). Most paleontologic studies dealing with late Paleozoic faunas focus on the extremely rich record of the Northern Hemisphere and, in addition, those on Gondwanic faunas have for the most part dealt with biostratigraphic rather than paleoecologic aspects. Interestingly, upper Paleozoic rocks of Gondwana commonly contain distinctive trace-fossil assemblages that are commonly preserved in postglacial deposits and, more rarely, in fine-grained intervals interbedded with or underlying glacial diamictites. Therefore, ichnologic evidence provides a proxy for use in reconstructing ecosystems associated to the Gondwana glaciation in rocks that are commonly devoid of body fossils. These peri-Gondwanic ichnofaunas are characterized by the dominance of simple grazing trails and arthropod trackways (Buatois et al., 2006). Their paleoecological and paleoenvironmental significance has largely been debated, with interpretations ranging from freshwater lacustrine to brackish-water deltaic or estuarine, and even fully marine (e.g., Savage, 1970, 1971; Anderson, 1970, 1975a, 1975b, 1976, 1981; Guerra-Sommer et al., 1984; Fernandes et al., 1987; Dias-Fabrício and Guerra-Sommer, 1989; Bercowski et al., 1990; Buatois and Mángano, 1992, 1993, 1995a, 1995b, 2003; Miller and Collinson, 1994; Peralta and Milana, 1999; Pazos, 2000, 2002a, 2002b; Trewin, 2000; Isbell et al., 2001; Nogueira and Netto, 2001a, 2001b; Balistieri et al., 2002, 2003a; Trewin et al., 2002; Buatois and del Papa, 2003; Gandini et al., 2007; Pazos et al., 2007). In an attempt to reconcile sedimentologic, ichnologic, and paleontologic data, it has recently been suggested that freshwater conditions—resulting from a strong discharge of freshwater due to melting of the ice masses—were prevalent in coastal areas during deglaciation (Buatois et al., 2006). Within this environmental scenario, glacial melting due to climatic amelioration led to the formation of freshwater water bodies that were physically connected with the open sea. In this setting, the simple dichotomy between marine and nonmarine settings is misleading. To complicate things further, some of the discussions surrounding the paleosalinity significance of these ichnofaunas are poorly constrained stratigraphically because trace-fossil assemblages coming from different stratigraphic levels are analyzed as a whole and paleoenvironmental interpretations are commonly uncritically extrapolated from one interval or locality to another. The aim of this paper is to further support the strong meltwater discharge hypothesis by evaluating the paleoecologic and paleoenvironmental significance of late Paleozoic ichnofaunas from eight different sedimentary basins along the Panthalassan margin of Gondwana. GLACIAL EVENTS AND POSTGLACIAL TRANSGRESSIONS: SEDIMENTOLOGIC AND ICHNOLOGIC SIGNATURES Three main glacial events are recognized in Gondwana, depending on paleolatitude: Devonian–Early Carboniferous, early Late Carboniferous, and Late Carboniferous–Early Perm-

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ian (López Gamundí, 1989; Limarino et al., 2002). Outcrops of the Devonian–Early Carboniferous glaciation are well exposed in the Titicaca Basin of Bolivia (e.g., Díaz Martínez et al., 1993), and the Amazonas and Parnaíba Basins of Brazil (Rocha-Campos, 1981a, 1981b). The early Late Carboniferous glacial event is well represented in the Andean basins of Argentina, including the Tarija, Calingasta-Uspallata, Paganzo, and San Rafael Basins (e.g., López Gamundí, 1997; del Papa and Martínez, 2001; Limarino et al., 2002; Marenssi et al., 2005), as well as the Sydney Basin of eastern Australia (e.g., Isbell et al., 2003a; Birgenheier et al., 2009). The Late Carboniferous– Early Permian glaciation is the most widespread and has been extensively documented in the Paraná Basin of Brazil, Paraguay, and Uruguay (e.g., Eyles et al., 1993; França, 1994; Vesely and Assine, 2006), the Karoo Basin of South Africa (e.g., Visser, 1990, 1997), the Kalahari Basin of Namibia, Botswana, and South Africa (e.g., Visser, 1997; Key et al., 1998), the Falkland Basin (Trewin et al., 2002), the Transantarctic Basin of Antarctica (e.g., Isbell et al., 1997), the Gondwana Master Basins of India (e.g., Wopfner and Casshyap, 1997), and the Sydney, Officer, Carnavon and Canning Basins of Australia (e.g., Veevers and Powell, 1987; O’Brien et al., 1998; Eyles and Eyles, 2000; Eyles et al., 2003). Our study is focused on the latter two of these three glacial events because no detailed ichnologic information is available for the Devonian–Early Carboniferous glacial event. Discussion of the ichnofaunas of the San Rafael, Karoo, Falkland, Transantarctic, and Sydney Basins is based on literature review and/or study of trace-fossil specimens. The Paganzo Basin (Western Argentina) The Upper Carboniferous–Upper Permian Paganzo Basin covers ~150,000 km2 of western Argentina (Fig. 1), and hosts sedimentary successions up to 3000 m thick. It is currently considered a foreland basin related to subduction of the Pacific plate beneath the western continental margin of Gondwana, which probably evolved to a rift system during the Permian (Ramos, 1988). The Protoprecordillera was a north-south-trending topographic high, albeit discontinuous, that separates mostly marine deposits of the Calingasta-Uspallata and Río Blanco Basins on the west from the mostly continental deposits of the Paganzo Basin on the east. The magmatic arc was located farther to the west, in present-day Chile. The Paganzo Basin was limited to the east and south by the Pampean and Pie de Palo topographic highs, respectively, whereas the Puna high represents its northern boundary. Sedimentation in the Paganzo Basin occurred in subbasins separated by internal basement highs. Overall, two major areas can be distinguished within the Paganzo Basin: an eastern zone dominated by continental environments, and a western one with increased marine influence (Limarino et al., 2002). Strata of the Paganzo Basin constitute the upper Paleozoic Paganzo Group (Fig. 2), although different nomenclatures have been established in various areas of the basin (Azcuy and Morelli, 1970). Glacial diamictites and postglacial fine-grained deposits

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occur in the lower Upper Carboniferous (Namurian–Westphalian) Guandacol Formation of the western area, and in the Agua Colorada Formation of the eastern area (Limarino et al., 2002). Farther to the east, the coeval Malanzán Formation records the same depositional cycle (Andreis et al., 1986; Buatois and Mángano, 1995a). Sedimentation began with coarse-grained alluvial fan deposits, braided fluvial deposits, and tillite that gave way rapidly to transgressive mudstone and sandstone (Limarino and Césari, 1988; Buatois and Mángano, 1995b; López Gamundí and Martínez, 2000; Limarino et al., 2002; Pazos, 2002a; Marenssi et al.,

2005). Deglaciation led to the establishment of large freshwater bodies to the east, and was accompanied by a marine incursion from the west that flooded valleys giving place to a series of fjords along the coastal area (Limarino et al., 2002; Kneller et al., 2004). These water bodies were filled during the subsequent regression mostly by deltaic progradation that eventually led to the establishment of fluvial systems across the area, represented in the overlying Upper Carboniferous–Lower Permian (Stephanian–Asselian) Tupe Formation and coeval units (Ottone and Azcuy, 1986; Limarino, 1988; Desjardins et al., 2009).

Ichnology of late Paleozoic postglacial transgressive deposits in Gondwana The ichnofauna of the Guandacol Formation has been analyzed in a number of studies, mostly in Cuesta de Huaco (Pazos, 2000) and Huerta de Huachi (Buatois and Mángano, 2003), San Juan Province. In Cuesta de Huaco, it has been subdivided into two main trace-fossil assemblages that yield insights into the ecological conditions during glacial retreat: the Didymaulichnus and the Mermia trace-fossil assemblages. Deposits bearing these two assemblages are separated by nonbioturbated black shale representing maximum flooding. The Didymaulichnus assemblage consists of monospecific suites of bilobate trails that historically have been assigned to Didymaulichnus lyelli, although the taxonomic affinities of these bilobate trace fossils are currently under revision. This assemblage occurs in a wide variety of lithologies, including mudstone, very fine- to fine-grained sandstone with current and combined-flow ripples, or granule conglomerate and very coarse-grained sandstone deposited from debris flows (Buatois et al., 2006). Dropstones are abundant in this interval. This assemblage is present in the lowermost strata of the transgressive interval, and has been found only in the western area. Lowdiversity suites of linguliformean brachiopods (cf. Oehlertella sp.) occur in one bed (Martínez, 1993). The ichnogenus Didymaulichnus occurs in both continental (e.g., Miller, 1986), and marine (e.g., Buatois and López Angriman, 1992) deposits. In the present context, the Didymaulichnus assemblage most likely represents brackish-water conditions in a stressful environment. Late Paleozoic brackish-water ichnofaunas are characterized by low-diversity (commonly monospecific) suites of very simple forms (Mángano and Buatois, 2004). The Mermia trace-fossil assemblage occurs in both Cuesta de Huaco and Huerta de Huachi (Figs. 3A–3F). It is relatively diverse and is dominated by nonspecialized grazing trails (Mermia carickensis, Gordia marina, Helminthopsis tenuis, Helminthoidichnites tenuis), simple feeding traces (Treptichnus pollardi, Circulichnis montanus), arthropod trackways (Diplichnites gouldi, Umfolozia isp., Maculichna carboniferus, Orchesteropus atavus), and fish trails (Undichna insolentia, Undichna britannica). Because these structures are preserved along bedding planes, recording emplacement in very shallow tiers, degree of bioturbation (as seen in cross section) is typically zero. The Mermia assemblage is present not only in the western area of the Paganzo Basin but also in the eastern region. In the Guandacol Formation, the Mermia assemblage occurs in parallel-laminated siltstone and current ripple cross-laminated and parallel-laminated very fine grained sandstone mostly deposited by delta-fed underflow currents and suspension fallout. This assemblage is present through the middle interval of the postglacial succession, representing early highstand systems tract deposition. This assemblage is typical of freshwater environments and represents an example of the Mermia ichnofacies. However, the presence of acritarchs in some beds, mostly in the western region, indicates a marine connection and the existence of intermittent periods of brackish-water conditions. Integration of palynologic and ichnologic data shows that acritarchs do not occur in beds containing trace fossils, which are commonly associated with ter-

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restrially derived palynomorphs (Buatois and Mángano, 2003). Body fossils are lacking in these deposits. The most likely environmental scenario for the Guandacol Formation is that of fjords filled by prograding deltaic deposits (Limarino et al., 2002). In the eastern region, the Mermia assemblage has been recorded in the coeval Agua Colorada Formation at Sierra de Narváez, Catamarca Province (Buatois and Mángano, 1993) (Fig. 4). High-diversity suites occur in parallel-laminated mudstone formed by suspension fallout and low-density turbidity currents, and in delta-fed, very fine grained sandstone and siltstone deposited from underflow currents. This ichnofauna includes Circulichnis montanus, Cochlichnus anguineus, Gordia marina, Gordia indianaensis, Helminthoidichnites tenuis, Helminthopsis tenuis, Mermia carickensis, Orchesteropus atavus, Rusophycus isp., Treptichnus pollardi, Undichna britannica, and U. insolentia, among other forms. A low-diversity suite, consisting of Gordia marina and Mermia carickensis, occurs at the top of tempestite and turbidite fine- to very fine grained sandstone, recording post-event colonization after major breaks in environmental conditions (Buatois and Mángano, 1995b). The palynologic assemblage consists only of terrestrially derived forms, and no acritarchs have been reported to date (Vergel et al., 1993). Integrated sedimentologic and ichnologic data together with the lack of marine indicators suggest deposition in large and deep lakes (Buatois and Mángano, 1994, 1995b), although deposition in the inner zone of fjords affected by strong meltwater discharge cannot be disregarded. Subsequently, a similar assemblage, although poorly preserved and less diverse, has been found in the Agua Colorada Formation at Bajo El Manzano (La Rioja Province). Trace fossils in this locality occur in parallel-laminated siltstone immediately above a dropstone-bearing siltstone interval. These fine-grained deposits are sandwiched between pebble conglomerate and very coarse to medium-grained sandstone that formed in low-sinuosity fluvial systems, and sandstone, siltstone, and conglomerate of fluvio-deltaic origin. Farther to the east, the coeval Malanzán Formation contains a low-diversity suite of grazing trails, recording colonization of turbidite sandstone (Buatois and Mángano, 1995a). Acritarchs have been recorded in the Malanzán succession, although not at the same levels as the trace fossils (Gutiérrez and Limarino, 2001). The San Rafael Basin (Western Argentina) The San Rafael Basin is a northwest-southeast trending late Paleozoic back-arc basin located in western Argentina (Espejo et al., 1996). It is limited to the south and southwest by the Neuquén Basin, and to the north and northeast by the Pampean Ranges. The Upper Carboniferous column is subdivided into the El Imperial and Agua Escondida Formations (Azcuy et al., 1999; Césari and Gutiérrez, 2000). The El Imperial Formation consists mostly of siltstone and sandstone, with diamictite toward the base of the succession. This unit has been subdivided into three main intervals (Arias and Azcuy, 1986). The lower interval consists of

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Figure 3. Trace fossils from the Guandacol Formation (Huerta de Huachi, Paganzo Basin, western Argentina). (A) Helminthoidichnites tenuis. (B) Helminthopsis tenuis. (C) Diplichnites gouldi showing poor preservation of individual appendage imprints. (D) Maculichna carboniferous. Preservational variety showing only one of the elements of the paired tracks. (E) Orchesteropus atavus. (F) Undichna britannica. Specimens housed at the Invertebrate Paleontology collection of the Instituto Miguel Lillo, Universidad Nacional de Tucumán, Argentina. All bars are 1 cm (after Buatois and Mángano, 2003).

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Figure 4. Trace fossils from the Agua Colorada Formation (Sierra de Narváez, Paganzo Basin, western Argentina). (A) Cochlichnus anguineus. (B) Helminthoidichnites tenuis. (C) Treptichnus pollardi. (D) Mermia carickensis. (E) Undichna insolentia. Specimens housed at the Invertebrate Paleontology collection of the Instituto Miguel Lillo, Universidad Nacional de Tucumán, Argentina. All bars are 1 cm (after Buatois and Mángano, 1993).

very coarse to fine-grained sandstone, siltstone, and pebble to cobble conglomerate deposited by turbidity currents (Arias and Azcuy, 1986). Pazos et al. (2007) mentioned evidence of wave action in some of the sandstone, indicating that at least part of the turbidite succession was emplaced in shallow water. The middle interval consists of diamictite and interbedded medium- to very fine grained sandstone and siltstone deposited in delta-front to delta-plain settings (Arias and Azcuy, 1986; Pazos et al., 2007). Fine-grained distal deposits commonly contain dropstones, whereas slumps occur in delta front deposits. The maximum flooding surface is contained within a black shale package. The upper interval consists of coarse- to fine-grained sandstone and siltstone recording deposition in anastomosed fluvial systems (Arias and Azcuy, 1986). The middle interval records the glacial

event, the postglacial transgression, and the subsequent highstand progradation (Pazos et al., 2007). The trace-fossil content of the glacial to postglacial event in the San Rafael Basin was recently analyzed by Pazos et al. (2007). These authors recognized three main trace-fossil assemblages in the El Imperial Formation. Their assemblage A occurs in the lower interval and is dominated by Diplopodichnus biformis and Diplichnites gouldi, with subordinate occurrences of Archaeonassa fossulata. Assemblage B is present in the lower part of the middle interval and contains simple feeding structures (Treptichnus pollardi), grazing trails (Helminthoidichnites tenuis, Mermia carickensis, Gordia marina, Cochlichnus anguineus), arthropod trails and trackways (Diplopodichnus biformis, Maculichna carboniferus), and fish trails (Undichna isp.). This assemblage

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favors comparison with the Mermia assemblage of the Guandacol and Agua Colorada Formations. Assemblage C occurs in the black shale package of the middle interval and consists only of Didymaulichnus lyelli and Diplopodichnus biformis. This assemblage is similar to the Didymaulichnus assemblage of the Guandacol Formation. Terrestrially derived palynomorphs occur throughout the whole succession (García, 1995). Rare acritarchs (Gorgonisphaeridium spp.) have been recorded in only one sample from the middle interval of El Imperial Formation (Pazos et al., 2007). Intense fluvial runoff into fjord environments and sporadic marine connections were indicated for the El Imperial Formation (Pazos et al., 2007). The Tarija Basin (Northwestern Argentina) The Tarija Basin extends ~1000 km from southeastern Peru in the north, through the Subandean region of Bolivia, and into northwestern Argentina toward the south (Fig. 1). It is considered an asymmetric intracratonic basin bounded by the Guapore Craton to the north and east, the Puna arch to the south and west, and the Michicola arch to the south and east (Eyles et al., 1995; Azcuy and di Pasquo, 1999; Starck and del Papa, 2006). In northwest Argentina, the Tarija Basin stratigraphic succession is subdivided into three groups: the Macharetí, Mandiyutí, and Cuevo. The lower Upper Carboniferous (Namurian–Westphalian) Macharetí Group contains rocks of the Gondwana glaciations, and is in turn subdivided into the Tupambi, Itacuamí, and Tarija Formations (Azcuy and di Pasquo, 1999). The Tupambi Formation is dominated by sandstone and represents sedimentation in periglacial environments characterized by estuarine and deltaic deposits filling incised valleys (López Gamundí, 1986; Starck et al., 1993; Starck and del Papa, 2006). The Itacuamí Formation consists of parallel-laminated mudstone with dropstones recording deposition in periglacial settings (Starck et al., 1993; del Papa and Martínez, 2001; Starck and del Papa, 2006). The Tarija Formation consists of glacial diamictite, medium- to fine-grained sandstone with current and combined-flow ripples, and parallel-laminated mudstone with dropstones (del Papa and Martínez, 2001; Starck and del Papa, 2006). Buatois and del Papa (2003) documented trace fossils in the uppermost part of the Itacuamí Formation, and in the lower interval of the Tarija Formation (Fig. 2). The Itacuamí ichnofauna consists of the locomotion trail Diplopodichnus biformis, locally intergrading with the trackway Diplichnites gouldi (Diplopodichnus-Diplichnites assemblage) (Fig. 5B). It occurs in suspension-fallout, parallel-laminated mudstone underlying diamictite. This assemblage is assigned to the Scoyenia ichnofacies. Overall, the ichnofauna from the Tarija Formation (Mermia assemblage) is dominated by tiny nonspecialized grazing trails (Mermia carickensis, Gordia marina, Helminthopsis tenuis, Helminthoidichnites tenuis), with subordinate locomotion traces (Cochlichnus isp., Diplopodichnus biformis, Diplichnites gouldi) (Fig. 5A, 5C, and 5D). The Tarija ichnofauna is present in thin

layers of very fine grained silty sandstone and siltstone, locally with rippled tops and convolute lamination, interpreted as the product of underflow currents, and thin mudstone partings covering medium- to fine-grained sandstone with wave ripples, wave ripple-cross lamination, and microhummocky cross-stratification, deposited by oscillatory flows. These deposits overlie diamictite and sandy turbidites, and underlie a thick diamictite package. The ichnofauna from the Tarija Formation is assigned to the Mermia ichnofacies. Taxonomic composition, nature of preservation, and overall features of the ichnofauna (e.g., presence of very simple grazing patterns of epifaunal animals and absence of trace fossils of infaunal organisms and typical ichnotaxa of marine environments) suggest freshwater conditions during accumulation of the ichnofossil-bearing deposits. This is consistent with the absence of marine body fossils and acritarchs in the unit. These conditions may have occurred in lakes or fjords affected by a strong discharge of fresh water due to melting of the ice masses. The Paraná Basin (Southern Brazil) The Paraná Basin is a widespread intracratonic depression that covered the entire southern portion of Brazil, as well as southeastern Paraguay, northeastern Argentina, and northern Uruguay (Fig. 1). It was formed during the Ordovician, being completely filled by the Late Cretaceous. During the Late Carboniferous– Early Permian, the landscape of the Paraná Basin area contained multi-lobed glacial fronts, forming fjord-like incised valleys opening toward a shallow epicontinental sea (Santos, 1987; Santos et al., 1996). The dominantly glaciomarine deposits of the Rio do Sul Formation represent the last glacial episode in the Paraná Basin, and the maximum flooding event related to the Gondwana deglaciation. Its eastern counterpart is represented by the Dwyka Series deposits, revealing a southeastern connection between the Paraná and the Karoo Basins (South Africa) during this time (Santos, 1987). Its southern counterpart is represented by the San Gregorio Formation from the North Uruguayan Basin (Goso, 1995). Six second-order sequences were recognized in the Paraná Basin (Milani, 1997). The Gondwana I sequence contains the glaciogenic deposits of the Upper Carboniferous–Lower Permian (Stephanian–Sakmarian) Itararé Group. In outcrop, the Itararé Group is subdivided into four formations: the Aquidauana, Campo do Tenente, Mafra, and Rio do Sul Formations (Schneider et al., 1974). Fluvial fine- to coarse-grained sandstone and glaciomarine diamictite are the main facies associations in the Aquidauana Formation, whereas glaciolacustrine massive siltstone and shale, and glaciofluvial diamictite, sandstone, and conglomerate dominate in the Campo do Tenente Formation (Castro et al., 1994). Striated pavements and striated and faceted clasts are common. The basin depocenter at this time was located in the midsouthern portion of Brazil (central and northern Santa Catarina State to south of Paraná State), where the thickest successions of the Itararé Group are well exposed. The Mafra and the Rio do

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Figure 5. Trace fossils from the Itacuamí, and Tarija Formations (Tarija Basin, northwest Argentina). (A) Diplopodichnus biformis. Negative epirelief in mudstone partings blanketing sandstone with wave ripples. Aguas Blancas. Tarija Formation. (B) Diplichnites gouldi (white arrow), and Diplopodichnus biformis (gray arrow). Negative epirelief in laminated mudstone. Arroyo Iquira. Itacuamí Formation. PIL 14843. (C) Mermia carickensis preserved in mudstone parting blanketing sandstone with wave ripples. Aguas Blancas. Tarija Formation. PIL 14854. (D) Helminthopsis tenuis (H), Diplopodichnus biformis transitional with Diplichnites gouldi (D), and Gordia marina (G). Positive hyporelief in sandstone with wave ripples. Aguas Blancas. Tarija Formation. Specimens housed at the Invertebrate Paleontology collection of the Instituto Miguel Lillo, Universidad Nacional de Tucumán, Argentina. All bars are 1 cm (after Buatois and del Papa, 2003).

Sul Formations are composed chiefly of deposits that originated during deglaciation (Fig. 2). The lower interval of the Mafra Formation consists of a glacio-deltaic succession that is replaced upward by a tide-influenced shallow-marine succession, consisting of medium- to large-scale, trough cross-stratified sandstone representing tidal bars (Schneider et al., 1974). The middle interval consists of very thinly bedded silty-muddy rhythmites filling an incised valley. Dropstones are abundant, ranging from granules (more common) to boulders. The upper interval of the Mafra Formation consists of thinly interbedded fine- to very fine grained sandstone and siltstone, and fine- to medium-grained sandstone with trough cross-stratification, herringbone cross-stratification,

and possibly hummocky cross-stratification. Matrix-supported diamictite with faceted clasts occurs throughout the whole succession. Widespread marine fossiliferous shale recording maximum flooding characterizes the base of the overlying Rio do Sul Formation, which consists mostly of fine-grained heterolithic deposits; diamictites are less abundant than in the Mafra Formation. Dropstones of different sizes are present. The ichnology of the Itararé Group has been analyzed in numerous papers (Guerra-Sommer et al., 1984; Fernandes et al., 1987; Dias-Fabrício and Guerra-Sommer, 1989; Nogueira and Netto, 2001a, 2001b; Balistieri and Netto, 2002; Balistieri et al., 2002, 2003a; Lermen, 2006; Gandini et al., 2007; Netto et al.,

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2009). The rhythmites of the Mafra Formation are well exposed in northern Santa Catarina State, where they make up part of the middle interval of the unit and contain two distinctive trace-fossil assemblages (Balistieri et al., 2002, 2003a). The first assemblage (Diplopodichnus-Diplichnites assemblage) occurs all throughout the rhythmite package, and consists exclusively of arthropod trackways (Diplichnites gouldi) (Fig. 6A) and trails (Diplopodichnus biformis) commonly associated with wrinkle marks. Both ichnotaxa are commonly intergradational, reflecting production by the same tracemaker, most likely a myriapod, but in a substrate ranging from soft (Diplopodichnus biformis) to slightly firm (Diplichnites gouldi) (Johnson et al., 1994; Keighley and Pickerill, 1996; Buatois et al., 1998; Paz et al., 2002, 2003, 2004; Balistieri, 2003; Balistieri et al., 2002, 2003a; Netto et al., 2009). Balistieri et al. (2002, 2003a) interpreted this trace-fossil assemblage as recording arthropod locomotion during periods of subaerial exposure of the substrate, most likely in mud flats flanking the fjord valley. Dryness is important in glacial climates, and promotes constant water evaporation and sublimation, especially during periods of colder temperatures when shallow-water bodies usually freeze. In contrast, during periods of higher temperatures, these shallow-water bodies defrost, supplying sediment to outwash plains. The second assemblage (Mermia assemblage) occurs toward the upper part of the rhythmite package and is dominated by nonspecialized grazing trails (e.g., Cochlichnus anguineus, Gordia arcuata, Gordia marina, Helminthoidichnites tenuis) and feeding traces (Hormosiroidea meandrica, Treptichnus pollardi) (Fig. 6B). Accessory components include arthropod locomotion (Cruziana cf. problematica) and resting traces (Rusophycus cf. carbonarius, Gluckstadtella isp.) and fish trails (Undichna consulca). Shallow grazing and feeding are the main ethological

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behaviors, revealing subaqueous colonization of soft substrates. The behavioral characteristics of the assemblage, the absence of truly marine ichnotaxa, the presence of tiny trails, and the dominance of simple feeding strategies are consistent with characteristics of the Mermia ichnofacies, suggesting colonization by a freshwater fauna (Balistieri, 2003; Balistieri et al., 2002, 2003a; Netto et al., 2009). Notably, in the thicker-bedded rhythmites, both trace-fossil assemblages coexist, representing distinct suites, with elements of the Mermia ichnofacies crosscut by the terrestrial suite, suggesting that the ponds periodically dried up (see Buatois and Mángano, 2004, 2007). The upper interval of the Mafra Formation contains diamictite and fine-grained heterolithic deposits that occur above the marine shale interval. The heterolithic deposits consist of interbedded fine- to very fine grained sandstone, and of siltstone with flaser and wavy bedding. The heterolithic facies contains ?Arenicolites isp., Chondrites isp., Diplocraterion isp., Palaeophycus isp., Planolites isp., Rhizocorallium isp., and Thalassinoides isp. (Balistieri, 2003). This trace-fossil assemblage occurs at several stratigraphic intervals within the Mafra Formation and extends into the overlying Rio do Sul Formation, invariably associated with heterolithic tide-influenced facies. Most ichnotaxa in this assemblage are facies-crossing trophic generalists, a strong evidence of opportunistic behavior (Ekdale, 1985; Bromley, 1996). These characteristics suggest substrate colonization by a brackish-water fauna, producing an impoverished mixed Cruziana-Skolithos ichnofacies (Pemberton and Wightman, 1992; MacEachern and Pemberton, 1994; Pemberton et al., 2001; Netto and Rossetti, 2003; Buatois et al., 2005; MacEachern and Gingras, 2007). At the base of Rio do Sul Formation, heterolithic deposits are punctuated by massive siltstone that contains bivalves, articulated brachiopods (Rocha-Campos, 1970), and a Glossifungites

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Figure 6. Trace fossils from rhythmites of the Mafra Formation (northern Santa Catarina State, Paraná Basin, southern Brazil). (A) Diplichnites gouldi. (B) Hormosiroidea meandrica. Specimens housed at the Paleontology collection of Centro Paleontológico de Mafra, Brazil. All bars are 1 cm (after Balistieri et al., 2002, 2003a).

Ichnology of late Paleozoic postglacial transgressive deposits in Gondwana suite (Balistieri and Netto, 2002) (Fig. 7). This suite is dominated by Thalassinoides isp., and also contains Diplocraterion isp., Palaeophycus isp., P. striatus, ?Rhizocorallium isp., and Gyrolithes-like burrows (see Netto et al., 2007), revealing burrowing in compacted, firm mud (Balistieri and Netto, 2002). The substrate-controlled Glossifungites suite delineates a firmground surface formed by erosional exhumation during transgression (MacEachern et al., 1992; Pemberton et al., 2001). The transgressive trend is further revealed by the presence of fossiliferous black shale (Lontras Shale) overlying this interval. This shale contains linguliformean brachiopods (Lingula sp. and Orbiculoidea guaraunensis), articulated brachiopods (Anthraconeilo sp., Barroisella imbituvensis, Chonetes riograndensis, Chonetes rionegrense, Crurithyris aff. planoconvexa, Crurithyris roxoi, and Langella imbituvensis) bivalves (Leda woodworthi and Nuculana woodworthi), paleoniscid and celacanthiid scales, shark teeth, arenaceous forams, and sponge spicules (Ruedmann, 1929; Oliveira, 1930; Mendes, 1952; Mezzalira, 1956; RochaCampos, 1970). Terrestrial elements in this shale interval include blattoid wings and terminal stems of gymnosperms (Pinto and Sedor, 2000). The shale interval indicates flooding of marginalmarine areas during the transgressive peak. The lower interval of the Rio do Sul Formation is also exposed in northern Santa Catarina State. These deposits contain abundant Protovirgularia isp., arthropod trackways (Diplichnites gouldi, Diplopodichnus biformis, Maculichna varia, Umfolozia sinuosa), common shallow burrows and trails (Cochlichnus anguineus, Hormosiroidea meandrica, Treptichnus isp.), and rare

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arthropod resting traces (Gluckstadella cooperi) (Gandini et al., 2007) (Fig. 8). Balistieri et al. (2002) also recorded chevronate trails in rhythmites of the base of the Rio do Sul Formation, assigned to the ichnogenus Protovirgularia, and noted that structures previously assigned to Gyrochorte isp. by Guerra-Sommer et al. (1984) and Marques-Toigo et al. (1989) most likely represent Protovirgularia. Regardless of the ichnotaxonomic assignment, these structures are most likely produced by arthropods rather than bivalves. The two arms of the V-shaped structures are not perfectly articulated at the axial area, indicating that they are not the result of the impression of a split muscular foot, but rather reflect the active and passive phases of an appendage stroke. The upper interval of the Rio do Sul Formation, which crops out in central Santa Catarina State, contains arthropod trackways (Diplichnites isp., Umfolozia isp.), trails (Cruziana isp., Diplopodichnus isp.), and resting trace fossils (Gluckstadtella isp., Rusophycus isp.), together with grazing trails (Helminthoidichnites isp.) (Nogueira and Netto, 2001a, 2001b). The Rio do Sul assemblage is similar to the assemblages recorded in the rhythmites of the Mafra Formation, albeit with a lower participation of grazing trails. The sporadic presence of marine plankton indicates periodic connection with the sea. In the southernmost portion of the Paraná Basin (Rio Grande do Sul State), the glaciogenic deposits of the Itararé Group have not been divided into formations (Schneider et al., 1974). These consist of very thin-bedded silty-muddy rhythmites overlying thickbedded diamictite and interbedded with fine-grained sandstone, interbedded fine-grained sandstone and siltstone, and black shale

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Figure 7. Trace fossils from a ravinement surface at the top of the Mafra Formation (northern Santa Catarina State, Paraná Basin, southern Brazil). (A) Diplocraterion isp. (B) Thalassinoides isp. Specimens not collected. All bars are 1 cm (after Balistieri and Netto, 2002).

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Figure 8. Trace fossils from the middle interval of Mafra Formation, and lower interval of the Rio do Sul Formation (northern Santa Catarina State, Paraná Basin, southern Brazil). (A) Gluckstadella cooperi (Gc) overprinted to Protovirgularia isp., forming a palimpsest surface (P). (B) Treptichnus isp. (C) Diplopodichnus biformis (Db), Maculichna varia (Mv), Hormosiroidea meandrica (Hm), and Diplichnites gouldi (Dg). (D) Palimpsest surface with Umfolozia sinuosa (Us), intergradational Diplopodichnus biformis (Db) and Diplichnites gouldi (Dg), Maculichna varia (Mv), and Hormosiroidea meandrica (Hm). Specimens housed at the Paleontology collection of the Centro Paleontológico de Mafra, and at the Invertebrate Paleontology collection of Universidade do Vale do Rio dos Sinos, Brazil. All bars are 1 cm (after Gandini et al., 2007).

(Correa da Silva, 1978). Dropstones are common in fine-grained deposits, and diamictites contain striated and faceted clasts. Striated pavements suggest long periods of nondeposition, possibly related to ice cap migration over highlands formed by basement rocks (Tomazelli and Soliani, 1982). The black shale (locally known as Budó Shale) contains a mix of marine and nonmarine fossils, composed of bivalves (Aviculopecten cambahyensis), brachiopods (Langella imbituvensis, Orbiculoidea maricaensis), marine scolecodonts (Nereidavus moreirai, Nereidavus beetleae, Ildraites langei, Arabellites almeidai, Arabellites santosi), ostracods, sponge spicules, fish remains (Elasmobranchii, paleonisciform and cladontiform teeth, coelacanthid and paleoniscid scales), paraplecopteran insects (Narkemina rochacamposi), a Glossopteris flora, and algal remains (Dolianiti, 1945;

Pinto, 1947, 1949, 1955, 2000; Martins, 1948, 1951; Martins and Sena Sobrinho, 1950; Pinto and Purper, 2000; Richter, 2000). This fossil assemblage is, in part, similar to that preserved in the Lontras Shale (Santa Catarina State) and the Passinho Shale (Paraná State). Trace fossils from the Itararé Group rhythmites of Rio Grande do Sul State were originally reported by Guerra-Sommer et al. (1984) and Dias-Fabrício and Guerra-Sommer (1989). A recent ichnotaxonomic review of this material (Lermen, 2006) reveals the presence of two distinct paleoichnocoenoses, one dominated by arthropod trackways (Fig. 9) and preserved in thin silty-muddy rhythmites, and the other one recorded by a Chondrites-Planolites-Palaeophycus composite ichnofabric, occurring in heterolithic fine-grained sandstone and siltstone. Shallow burrows, as

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Figure 9. Trace fossils from the Itararé Group rhythmites (Rio Grande do Sul State, Paraná Basin, southern Brazil). (A) Cruziana problematica. (B) Maculichna varia. (C) Kouphichnium isp. (D) Protichnites isp. Specimens housed at the Paleobotany Paleontology collection of Universidade Federal do Rio Grande do Sul, Brazil. All bars are 1 cm (after Lermen, 2006).

well as arthropod resting traces and intrastratal trails, are subordinate elements in the rhythmite paleoichnocoenosis. Trace-fossil information suggests a marginal-marine, possibly fjord-like depositional setting, similar to that represented by the Mafra Formation deposits. The rhythmite ichnocoenosis has a strong freshwater signature, whereas the Chondrites-PlanolitesPalaeophycus composite ichnofabric most likely formed under brackish-water conditions. The dominance of arthropod trackways in the rhythmite ichnocoenosis suggests affinities with the Scoyenia ichnofacies, whereas the heterolithic-facies ichnocoenosis represents an impoverished Cruziana ichnofacies. The trace-fossil content of the Itararé Group in Rio Grande do Sul State is similar to that preserved in rhythmites of the Rio do Sul Formation in northern and central Santa Catarina State, further supporting the hypothesis that these deposits represent the last deglaciation events in the Paraná Basin. The Paraná Basin extends toward the south into Uruguay (the so-called North Uruguayan Basin), occupying an area of ~90,000 km2. This basin was filled mostly by siliciclastic Paleozoic and Mesozoic deposits (Goso, 1995). Four major sedimentary sequences have been recognized in this region, ranging from Middle Devonian to Paleogene. The glacial-influenced deposits of the San Gregorio Formation correspond to the initial deposition of the Permian–Triassic sequence represented by the Cerro Largo Group (Preciozzi et al., 1988). The San Gregorio Formation (equivalent of the Itararé Group and the middle interval of the Ecca Group) is regionally exten-

sive and reflects deposition controlled by basement highs (Goso, 1995). This unit contains a relatively diverse marine fauna, including orthocone cephalopods (Dolorthoceras chubutensis), goniatites (Eoasianites (Glaphyrites) rionegrensis), radiolarians, sponges (Itararella gracilis, Microhemidis ortmanni), fish remains (Coccocephalichthys tesselatus, Carbonilepsis uruguayensis, Daphnaechelus formosus, Elonichthys macropercularis, Gondwanaichthys maximus, Itaratichthys microphtalmus, Mesonichthys antipodeus, Rhadinichthys rioniger), and arthropods (Martínez Machiavello, 1963; Closs, 1967, 1969; Kling and Reif, 1969; Ybert and Marques-Toigo, 1970; Marques-Toigo, 1970, 1973a, 1973b, 1974; Beltan, 1978, 1981; Mones and Figueiras, 1980; da Silva, 1985; Beri, 1987, 1991; Beri and Daners, 1994). The precise age of the San Gregorio Formation has been debated. The cephalopod and fish remains suggest a Pennsylvanian age whereas the palynomorphs indicate Artinskian-Sakmarian. The San Gregorio Formation contains an ichnofauna that includes typical marine ichnogenera, such as Phycosiphon, Chondrites, Rhizocorallium, Thalassinoides, and Paleobullia, accompanied by some facies-crossing forms (e.g., Palaeophycus, Diplocraterion, Taenidium, Cruziana, Gordia, Planolites, Protichnites) (Netto and Goso, 1998; Balistieri et al., 2003b). This ichnofauna is similar to the mixed Skolithos-Cruziana association recorded in the heterolithic deposits of the lower interval of the Rio do Sul Formation in northern Santa Catarina State (southern Brazil). The low to moderate intensity of bioturbation and ichnodiversity, and the dominance of trace fossils

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made by trophic generalists, suggests brackish-water conditions in marginal-marine environments. This interpretation is consistent with those proposed by da Silva (1985), who regarded these glaciogenic deposits as estuarine on the basis of palynomorphs, and by Goso (1995), who suggested that the absence of marine plankton resulted from meltwater discharge. However, the abundance of ichnofabrics dominated by Chondrites and Phycosiphon points to periods of more marine influence, more typical of prodeltaic settings. The Karoo Basin (South Africa) The Karoo Basin is a retro-arc foreland basin located north of the Cape Fold Belt that contains up to 8 km of strata (Rubidge et al., 2000). It represents the remnant of a more extensive roughly north-south-trending basin along part of the Gondwanic paleo-Pacific margin (Visser, 1987). Reconstructions suggest that the most important ice-spreading centers were located over south-central Africa (Visser, 1997; Catuneanu, 2004). The Upper Paleozoic column of the Karoo Basin is included in the Karoo Supergroup. The Upper Carboniferous–Upper Permian succession has been subdivided into three groups, Dwyka, Ecca, and Beaufort (Visser, 1983, 1995, 1996, 1997; Catuneanu et al., 1998). The Upper Carboniferous–Lower Permian (Stephanian– Asselian) Dwyka Group consists of glacial diamictite separated by varve shale and siltstone, representing glacial and interstadial phases, respectively (Visser, 1983, 1997). The stratigraphy of the Lower–Middle Permian (Asselian– Capitanian) Ecca Group varies from section to section (Catuneanu et al., 2005). The lower Ecca includes three formations, Prince Albert, Whitehill, and Collingham. The Prince Albert Formation consists of mudstone with chert, carbonate, and phosphate nodules, recording syn- to postglacial suspension fallout and episodic deposition from turbidity currents and mud flows (Visser and Young, 1990). The Whitehill Formation consists of black organic-rich shale, representing deposition under anoxic conditions (Visser, 1992; Johnson et al., 1997). The Collingham Formation consists of chert, mudstone, ash, and sandstone, recording suspension fallout alternating with low-density turbidity currents in a basin floor (Viljoen, 1994). Toward the foreland, the upper Ecca is subdivided into the Tierberg, Skoorsteenberg, Kookfontein, and Waterford Formations (Wickens, 1996; Catuneanu et al., 2005). The Tierberg Formation consists of dark basinal shale, whereas the Skoorsteenberg Formation records five major sandrich turbidite systems separated by fine-grained intervals (Scott et al., 2000; Johnson et al., 2001). The Kookfontein Formation consists of thick sandstone units separated by thin mudstone intervals, mostly recording deltaic progradation (Wickens, 1996). The Waterford Formation consists of sandstone and siltstone, and records shallow-marine to paralic deposition (Rubidge et al., 2000; Catuneanu et al., 2005). The Upper Permian (Wuchiapingian– Changhsingian) to Middle Triassic Beaufort Group consists of interbedded mudstone and sandstone, recording deltaic, fluvial, and lacustrine environments (van Dijk et al., 1978).

Descriptions of the Dwyka and Ecca ichnofaunas have been provided by Savage (1970, 1971) and Anderson (1970, 1975a, 1975b, 1976, 1981). The Dwyka ichnofauna consists of arthropod trackways (Umfolozia sinuosa, U. longula, Maculichna varia, Diplichnites isp.) and resting traces (Gluckstadtella cooperi, Kingella natalensis), and fish trails (Undichna insolentia, U. simplicitas, U. bina, Undichna isp.). At least part of the material originally referred to Gyrochorte isp. by Savage (1971) may be included in Protovirgularia, although a bivalve affinity is questionable, and the structures were most likely produced by arthropods, as in the case of the specimens from the Paraná Basin. The nature of the tracemaker is significant because bivalve-generated chevronate structures are produced by cleft-foot bivalves, which are exclusively marine. Most of the trace fossils were recorded in a single quarry near Swart Umfolozi in northern Natal (Fig. 10), but other localities in western South Africa and southern Namibia have provided specimens, as well as a core in Orange Free State. The freshwater affinities of the Dwyka ichnofauna have been noted by almost all workers in the area (Anderson, 1970, 1975a, 1976, 1981; Savage, 1970, 1971). This is consistent with the restricted distribution of marine body fossils during Dwyka times (McLachlan and Anderson, 1973). Cephalopods and nuculanid bivalves have been recorded in the western part of the basin, whereas acritarchs are known from boreholes in the south. The Dwyka ichnofauna is virtually identical to the one described from the Rio do Sul Formation rhythmites in Brazil. The Ecca ichnofauna is more variable. Arthropod trackways (Umfolozia sinuosa, U. longula) and fish trails (Undichna insolentia, U. simplicitas, U. bina) have been described from a number of localities both in the southern “turbidite” facies and in the basin center facies (Anderson, 1970, 1976, 1981). Also, Kouphichnium isp. was recorded at the top of the Whitehill Formation, and in the upper Ecca close to the boundary with the Beaufort Group (Anderson, 1975b). Marine ichnotaxa (e.g., Chondrites, Lorenzinia, Lophoctenium) have been mentioned (but not illustrated) by Johnson et al. (2001) in turbidites from the Skoorsteenberg Formation in the southwest portion of the basin. Marine trace-fossil assemblages containing Siphonichnus eccaensis, Spirodesmos archimedeus, Rhizocorallium isp., Skolithos isp., and Diplocraterion isp. have been documented in deltaic deposits of the Middle Ecca in Natal (Hobday and Tavener-Smith, 1975; Stanistreet et al., 1980; Turner et al., 1981; Mason et al., 1983; Tavener-Smith and Mason, 1985). Other ichnotaxa described from the Ecca need to be reevaluated. Marine body fossils are rare in the Ecca. The most abundant marine fauna occurs in the Whitehill Formation and consists of the reptile Mesosaurus, the fish Palaeoniscus, and the crustacean Notocaris tapscotti (Smith et al., 1993; Johnson et al., 1997). Cephalopods have also been described in middle Ecca beds in Natal (Rilett, 1963; Teichert and Rilett, 1974), and bivalves are known from the Waterford Formation in southern Cape Province (Cooper and Kensley, 1984; Rubidge et al., 2000). Teichert and Rilett (1974) noted the presence of a bed containing a marine fauna 1.5 m above a bed with freshwater bivalves. Integration of ichnologic information

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within a sedimentologic, paleontologic, and stratigraphic framework is essential for unraveling paleosalinity conditions and the complex depositional history of the Ecca Group. Late Paleozoic deposits extend toward Namibia in the intracratonic Kalahari Basin. Paleontologic evidence from the Dwyka and Ecca Groups in southern Namibia indicates marine conditions (Martin and Wilczewski, 1970). Mollusc shells (e.g., Eurydesma mytiloides), bryozoan colonies, cephalopods, echinoderms, fishes (Namaichthys schroederi), algal mats, and algal-serpulid buildups occur in the Dwyka Group in Namibia (Dickins, 1961; Gardiner, 1962; Martin and Wilczewski, 1970; Grill, 1997; Visser, 1983). The trace fossils recorded in this unit are Rhizocorallium irregulare, Rosselia isp., and Planolites isp., indicating marine conditions (Grill, 1997). Also, a marine ichnofauna (Siphonichnus isp., Arenicolites isp.) has been documented by this author in shoreface deposits of the Auob Sandstone Member of the Prince Alpert Formation (Ecca Group). The Falkland Basin (Falkland Islands) The late Paleozoic Falkland Basin (Fig. 1) is thought to represent a “missing” southeast corner of the Karoo Basin after a 180° rotation of the Falkland block (Adie, 1952; Mitchell et al., 1986; Taylor and Shaw, 1989; Marshall, 1994; Trewin et al., 2002). Upper Paleozoic rocks are subdivided into the Bluff Cove, Fitzroy Tillite, Port Sussex, Brenton Loch, and Bay of Harbours Formations (Trewin et al., 2002). The Upper Carboniferous–Lower Permian (StephanianSakmarian) Fitzroy Tillite Formation is dominated by diamictite with subordinate presence of sandstone and mudstone, representing glacially related subaqueous debris flow and esker deposits

Figure 10. Trace fossils from the Dwyka Group (Swart Umfolozi River, Karoo Basin, South Africa). Umfolozia sinuosa (Us) and Undichna isp. (U). Slab is housed at the British Museum of Natural History, London. Scale bar is 1 cm.

(Frakes and Crowell, 1967; Trewin et al., 2002). Trewin et al. (2002) recorded the presence of the arthropod trackway Umfolozia in a mudstone unit near the base of this formation. The Lower– Middle Permian (Artinskian-Capitanian) Port Sussex Formation consists of sandstone and mudstone with dropstones, representing the postglacial transgression, commonly under anoxic conditions (Trewin et al., 2002). Thin-bedded turbidites occur toward the top of the formation (Shepherds Brook Member). The Middle Permian (Capitanian) Brenton Loch Formation consists of massive turbidites with abundant dewatering structures and rhythmically laminated sandstone and mudstone couplets, most likely emplaced from density underflows. This unit represents deposition in turbidite systems and prograding deltas (Trewin et al., 2002). Trewin et al. (2002) documented the ichnology of this unit, recording a relatively high diversity of trace fossils in the middle member of the formation (Cantera Member). The Cantera Member ichnofauna consists of arthropod trackways (Umfolozia longula, Kouphichnium isp.), arthropod resting trace fossils, fish trails (Undichna bina, U. cf. insolentia), simple grazing trails (Helminthoidichnites, Cochlichnus, Spirodesmos), simple feeding trace fossils (Planolites, Treptichnus), and vertical burrows (Diplocraterion). A similar, although less diverse, ichnofauna occurs in the upper Saladero Member. Trewin et al. (2002) noted the remarkable similarities between the Falkland ichnofaunas and those from the Paganzo Basin of Argentina and the Karoo Basin of South Africa, stressing the absence of any marine indicator. Therefore, they interpreted the late Paleozoic succession as representing deposition in a freshwater lake. These authors also noticed that freshwater conditions in the Falkland Basin were consistent with the absence of marine fossils in the eastern portion of the Karoo Basin.

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The Transantarctic Basin (Antarctica) The Transantarctic Basin of the central Transantarctic Mountains (Fig. 1) contains a thick succession of Permian–Jurassic continental rocks (Barrett et al., 1986; Collinson et al., 1994). The Permian column is subdivided in the Pagoda, Mackellar, Fairchild, and Buckley Formations. The Upper Carboniferous–Lower Permian Pagoda Formation consists of diamictite and sandstone recording periglacial and glacial deposition (Barrett et al., 1986; Miller, 1989; Isbell et al., 2001). The Lower Permian Mackellar Formation consists of fine-grained deposits, essentially shale, siltstone, and thin-bedded sandstone, and records the postglacial transgression in a large and deep lake affected by underflow and turbidity currents (Miller and Collinson, 1994; Collinson et al., 2004; Miller and Isbell, this volume). The Lower–Middle Permian Fairchild Formation consists of sandstone formed in lowsinuosity fluvial systems that prograded into the lake, forming a braided outwash plain (Collinson et al., 1994; Miller and Isbell, this volume). The Middle–Upper Permian Buckley Formation consists of sandstone, siltstone, shale, and coal formed in highsinuosity fluvial systems (Knepprath et al., 2004). The ichnology of postglacial deposits of the Mackellar Formation has been discussed by Miller and Collinson (1994) and by Miller and Isbell (this volume). The ichnofauna consists of simple grazing trails (Cochlichnus, Mermia), simple feeding trace fossils (Planolites, Treptichnus), horizontal dwelling burrows (Palaeophycus), arthropod scratch marks, and bilobate trails, illustrating the freshwater Mermia ichnofacies. As in the case of other Gondwana ichnofaunas, vertical bioturbation is negligible and all ichnotaxa are parallel to the bedding plane (Miller and Isbell, this volume). Geochemical data (high C:S ratios) also support freshwater conditions (Miller and Collinson, 1994). The Sydney Basin (Eastern Australia) The Sydney Basin (Fig. 1), located in New South Wales, eastern Australia, covers 64,000 km2 both onshore and offshore, and includes up to 6000 m of Carboniferous–Triassic strata (Gostin and Herbert, 1973; Fielding et al., 2001). The basin is considered a foreland basin formed in connection with the New England Fold Belt, which developed as a volcanic arc during the Late Carboniferous (McPhie, 1987). In the northern area of the basin, the bulk of sediment fill is included, from base to top, in the Lower Permian Dalwood Group, the Lower to Upper Permian Maitland Group, and the Upper Permian Singleton Supergroup. However, several less well known formations (Currabubula Formation, Spion Kop Conglomerate, and Seaham Formation) occur below the Permian–Triassic column as a series of unconformitybounded units that include Upper Carboniferous glacial deposits of the second glacial episode (Isbell et al., 2003a; Birgenheier et al., 2009). Of these, the Seaham Formation of earliest Namurian to earliest Westphalian age overlies a polished and striated surface cut on the underlying Paterson Volcanics. It consists of diamictite,

tuff and thinly interbedded sandstone, and siltstone, and has been interpreted as deposited in fluvial and lacustrine environments adjacent to a volcanic arc (Campbell, 1969; Benson, 1981; Isbell et al., 2003a; Birgenheier et al., 2009). Detailed integrated sedimentologic and ichnologic studies are lacking, but an ichnofauna of arthropod trackways (Umfolozia sinuosa) and trails (Cruziana problematica) is present in a siltstone interval overlying glacial diamictite (Fig. 11). DISCUSSION Recent research in late Quaternary marine deposits associated with glacial meltwater has emphasized the role of hyperpycnal flows in connection with high freshwater discharges (e.g., Piper et al., 2007; Tripsanas et al., 2007). High freshwater discharges due to glacier melting and associated catastrophic outburst floods have been documented for the Baltic and Labrador Seas (Lord, 1990; Shaw and Lesemann, 2003). In addition, times of elevated concentration of suspended sediment promote hyperpycnal flows in modern fjords (Syvitski et al., 1987). Only exceptional discharges overcome the buoyancy effect of seawater in modern examples, but high-discharge hyperpycnal flows may have been extremely common during deglaciation. Also, because large discharges reduce the salinity of the water body, the likelihood of hyperpycnal flows is increased, providing a positive feedback. As a result, glacial melting led to the formation of freshwater bodies that were physically connected with the open sea. For example, the Holocene Yoldia Sea was freshwater in the northern Baltic Sea Basin due to a high input of meltwater during deglaciation during most of its history (Virtasalo et al., 2006). A wide variety of environmental stresses that affect benthic colonization may play a role in these glacially influenced environments. These include extreme dilution of salinity, high rates of sedimentation, variable degrees of substrate consolidation, oxygen-depleted conditions, high water turbidity, and intense storm activity (Feder and Keiser, 1980; Feder and Matheke, 1980; Farrow et al., 1983; Syvitski et al., 1987; O’Clair and Zimmerman, 1987). In addition, seasonal light restriction and floating ice masses contributing to ice-rafted debris rainfall may be important stress factors in polar areas. Integration of ichnologic evidence with sedimentologic, stratigraphic, and paleontologic data allows reconciling apparently inconsistent sets of data and sheds light on the nature of peri-Gondwanic ecosystems (Fig. 12). A complex paleogeography of fjords and deep, large coastal lakes is proposed for the Panthalassan margin of Gondwana. Although these peri-Gondwanic coasts were unique in some sense, comparisons with the modern coasts of southern Chile and Norway are suggested. All eight basins discussed in this paper contain glacial diamictites, postglacial fines, and distinctive ichnofaunas. Although the units differ in the degree of marine connection, the common theme in all is the presence of freshwater ichnofaunas in direct association with glacially influenced coasts affected by strong discharges of meltwater. At one extreme, the Itacuamí and Tarija Formations of the Tarija Basin, the Agua Colorada of the Paganzo Basin,

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Us Figure 11. Trace fossils from the Seaham Formation (Sydney Basin, eastern Australia). Umfolozia sinuosa (Us) and Cruziana problematica (Cp). Slab is housed at the Invertebrate Paleontology collection of Macquarie University, Sydney. Scale bar is 1 cm.

Us

Us

the Fitzroy Tillite and Brenton Loch Formations of the Falkland Basin, and the Mackellar Formation of the Transantarctic Basin contain only freshwater ichnofaunas, and no in situ marine palynomorphs or body fossils have been found. Accordingly, they may represent deposition either in freshwater lakes isolated from the sea or in fjords affected by strong meltwater discharge. The other end member is represented by the Mafra and Rio do Sul Formations of the Paraná Basin and the Ecca Group of the Karoo Basin, which contain intervals with relatively abundant marine faunas and ichnofaunas indicative of marine conditions (depauperate Cruziana and Glossifungites ichnofacies). Intermediate cases are represented by the Guandacol Formation of the Paganzo Basin, and the El Imperial Formation of the San Rafael Basin. The Guandacol Formation contains a single bed with linguliformean brachiopods near its base and discrete levels with acritarchs (Martínez, 1993). One level with acritarchs has been recorded in the El Imperial Formation (Pazos et al., 2007). Ichnologic information from late Paleozoic peri-Gondwanic settings indicates that freshwater conditions prevailed in coastal areas during deglaciation because of extreme discharge of freshwater due to melting of ice masses (Buatois et al., 2006). In some of these basins (e.g., Paganzo, Paraná), glacial melting led to the formation of freshwater bodies that were physically connected with the open sea. Recent reconstructions indicate that ice sheet volumes of late Paleozoic Gondwana were in the range of 49.1– 65.4 × 106 km3, although lower volumes have been estimated for many ice sheets (Isbell et al., 2003a). Complete ablation of a single ice sheet would have resulted in a 100 m sea-level rise with a concomitant reduction in coastal salinity. In these situations, establishment of a marine benthos is inhibited due to reduced salinity, allowing colonization by a freshwater biota. In addition, the dominance of horizontal feeding traces of deposit and detritus

Cp

feeders and the absence of vertical burrows of suspension feeders are consistent with high amounts of suspended fine-grained material (MacEachern et al., 2005). The fact that ichnofaunas occur throughout relatively thick intervals (e.g., ~30 m in the Guandacol Formation) suggests that freshwater conditions were temporally persistent in some of the analyzed basins (Buatois et al., 2006). Also, freshwater ichnofaunas can be traced for ~250 km along dip in the Paganzo Basin, reflecting remarkable seaward migration of the salinity barrier. Although some of these settings have been referred to as “brackish seas,” in fact they may be more appropriately called “freshwater seas” because of the dominance of freshwater conditions due to extensive melting during postglacial times. Recently, geochemical analyses were performed in some of the postglacial units of South Africa (Scheffler et al., 2006; Herbert and Compton, 2007). According to Herbert and Compton (2007), the texture, low δ18O values, and radiogenic Sr isotope ratios of early diagenetic calcite concretions of the Dwyka Group and phosphatic chert concretions of the Prince Albert Formation of the Ecca Group support an origin in glacial, freshwater sediments. In contrast, Scheffler et al. (2006) concluded from low CIA, Rb/K, V/Cr and TOC contents that the upper Dwyka Group was characterized by an alternation of freshwater and brackishwater conditions in the southern part of the basin. In addition, on the basis of Rb/K ratios, they noted that more marine conditions alternated with short-term brackish-water conditions during deposition of the lower Prince Albert Formation. There are relatively few studies documenting ichnofaunas in Quaternary glacial to postglacial settings. Most of these studies focus on Pleistocene glacial lakes (e.g., Gibbard and Stuart, 1974; Gibbard, 1977; Gibbard and Dreimanis, 1978; Walter and Suhr, 1998; Gaigalas and Uchman, 2004; Benner and Ridge,

M

D C

M

Un

C D Dp

G

Un

Ar

Rz

Th

Di

Pl

Di

Ch

Pa

- Marine palynomorphs - Diverse marine fauna in shale intervals overlying ravinement surfaces - e.g., Mafra Formation

- Terrestrially derived and marine palynomorphs - Marine Fauna - eg., Ecca Group, Mafra & Rio do Sul formations

Th

Depauperate Cruziana–Skolithos Ichnofacies

Glossifungites Ichnofacies

- Terrestrially derived palynomorphs - Local presence of acritarchs and/or linguliformeans in adjacent intervals - e.g., Guandacol and El Imperial formations

Figure 12. Reconstruction of postglacial coastal environments and associated ichnofaunas along the Panthalassan margin of Gondwana. M—Mermia, C—Cochlichnus, G—Gordia, Un—Undichna, D—Diplichnites, Dp—Diplopodichnus, U—Umfolozia, Ma—Maculichna, Ar—Arenicolites, Di—Diplocraterion, Th—Thalassinoides, Pa—Palaeophycus, Pl— Planolites, Ch—Chondrites, Rz—Rhizocorallium. Not all trace-fossil assemblages are present in every depositional system or basin. Changes in coastal physiography in response to sea-level changes, and degree of freshwater discharge result in ichnofaunal changes through time and, accordingly, some trace-fossil assemblages may be replaced by others, reflecting different paleosalinity levels. Although marine influence overall increases in a seaward direction, the presence of highly irregular shores with common lateral tributaries makes linear proximal-distal extrapolations an oversimplification at best (e.g., freshwater coastal lakes between embayed fjord areas may occur in a distal position with respect to adjacent marine tongues).

- Terrestrially derived palynomorphs - No marine fossils - e.g., Agua Colorada, Fritzroy Tillite & Mackellar formations

G

Ma

Mermia–Scoyenia Ichnofacies

U

Mermia–Scoyenia Ichnofacies

Ichnology of late Paleozoic postglacial transgressive deposits in Gondwana 2007; Benner et al., 2009; Uchman et al., 2009). Pleistocene glacial lake ichnofaunas are dominated by nonspecialized grazing trails, arthropod trackways, and, to a lesser extent, fish trails, closely resembling Gondwana fjord and lacustrine ichnofaunas. Brackish-water ichnofaunas, consisting of bivalve vertical burrows (Siphonichnus), U-shaped vertical burrows (e.g., Arenicolites, Diplocraterion), gravel-lined polychaete burrows (Diopatrichnus), and crustacean galleries (Thalassinoides), among other forms, have been documented in Cenozoic fjord deposits of Alaska (Eyles et al., 1992). Monospecific suites of Teichichnus occur in Holocene fjord deposits of Norway (Corner and Fjalstad, 1993). A recent study of Holocene cores from the Baltic Sea by Virtasalo et al. (2006) demonstrated a vertical change of ichnofaunas in connection to a shift from freshwater to brackishwater conditions. The trace-fossil assemblage of the Ancylus Lake consists of Palaeophycus and Arenicolites, representing an extremely impoverished ichnofauna. With the advent of the brackish-water Littorina Sea, this assemblage was replaced by an assemblage consisting of Planolites, Arenicolites, Lockeia, and Teichichnus. Although an increase in ichnodiversity occurs as a result of glacial melting, ichnodiversity levels remain very low as a result of brackish-water conditions during establishment of the Littorina Sea. This pattern of vertical distribution of trace fossils in response to a transition from freshwater to brackish water is similar to that observed in the Paganzo and Paraná Basins, albeit at different scales (Balistieri and Netto, 2002; Desjardins et al., this volume). Pazos et al. (2007) noted that most models for brackish environments are based on nonglacial successions and speculated that they may not be applicable to glacial settings. However, our review demonstrates that, although many other environmental controls obviously play a role, salinity is also of importance, and in fact brackish-water ichnofaunas are well known in late Paleozoic marginal-marine deposits of Gondwana and in Quaternary fjord deposits (see also Buatois et al., 2005). Interestingly, although coastal deposits not influenced by glaciation contain fully marine ichnofaunas in the transgressive maximum interval, transgressive and early highstand deposits in glacially influenced basins may contain freshwater ichnofaunas, as a result of extreme meltwater release. This should be taken into account if trace fossils are used in sequence-stratigraphic studies of Gondwana successions. We propose that in the case of Gondwana ichnofaunas the simple dichotomy between marine and nonmarine settings is misleading, because these peculiar assemblages should first be understood in terms of their paleoecologic significance, and subsequently placed within a larger paleoenvironmental context that takes into account the unusual depositional conditions associated with Gondwana glaciations and subsequent glacial retreat. This pattern indicates the existence of laterally persistent, albeit diachronous, peri-Gondwanan ichnofaunas that characterize melting of the late Paleozoic ice caps. Temporal recurrence of these ichnofaunas through the Late Carboniferous–Middle Permian indicates a common response of benthic faunas under similar ecological conditions during deglaciation events.

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ACKNOWLEDGMENTS Our ideas on ecologic aspects of the Gondwana glaciation grew out of a series of projects undertaken in many late Paleozoic basins. We would like to thank Cecilia del Papa, Oscar Limarino, and Ernesto Lavina for valuable discussions. Alfred Uchman and Ricardo Melchor provided valuable reviews of the manuscript. Patricio Desjardins prepared Figure 11 and Terri Graham assisted with the reference list. Financial support for this study was provided by the Brazilian National Council for Scientific and Technological Development (CNPq) Grants 474345/2003-3 and 304811/2004-1 awarded to Netto, and by the Canadian Natural Sciences and Engineering Research Council (NSERC) Discovery Grants 311727-05/08 and 311726-05/08 awarded to Mángano and Buatois, respectively. REFERENCES CITED Adie, R.J., 1952, The position of the Falkland islands in a reconstruction of Gondwanaland: Geological Magazine, v. 89, p. 401–410, doi: 10.1017/ S0016756800068102. Aitken, A.E., 1990, Fossilization potential of Arctic fjord and continental shelf benthic macrofaunas, in Dowdeswell, J.A., and Scourse, J.E., eds., Glacimarine Environments: Processes and Sediments: Geological Society [London] Special Publication 53, p. 155–176. Anderson, A., 1970, An analysis of supposed fish trails from interglacial sediments in the Dwyka Series, near Vryheid, Natal, in Proceedings, International Union of Geological Sciences Gondwana Symposium, 2nd, Pretoria, South Africa, p. 637–647. Anderson, A., 1975a, Turbidites and arthropod trackways in the Dwyka glacial deposits (Early Permian) of southern Africa: Transactions, Geological Society of South Africa, v. 78, p. 265–273. Anderson, A., 1975b, Limulid trackways in the Late Palaeozoic Ecca sediments and their palaeoenvironmental significance: South African Journal of Science, v. 71, p. 249–251. Anderson, A.M., 1976, Fish trails from the early Permian of South Africa: Palaeontology, v. 9, p. 397–409. Anderson, A., 1981, The Umfolozia arthropod trackways in the Permian Dwyka and Ecca Series of South Africa: Journal of Paleontology, v. 55, p. 84–108. Andreis, R., Leguizamon, R., and Archangelsky, S., 1986, El Paleovalle de Malanzán: Nuevos criterios para la estratigrafía del Neopaleozoico de la Sierra de Los Llanos, La Rioja, República Argentina: Boletín Academia Nacional de Ciencias, v. 57, p. 2–122. Angiolini, L., Balini, M., Garzanti, E., Nicora, A., and Tintori, A., 2003, Gondwanan deglaciation and opening of Neotethys: The Al Khlata and Saiwan formations of Interior Oman: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 196, p. 99–123, doi: 10.1016/S0031-0182(03)00315-8. Arias, W.E., and Azcuy, C.A., 1986, El Paleozoico superior del cañón del río Atuel, provincia de Mendoza: Revista de la Asociación Geológica Argentina, v. 41, p. 262–269. Azcuy, C.A., and Morelli, J.R., 1970, Geología de la comarca Paganzo-Amaná. El Grupo Paganzo: Formaciones que lo componen y sus relaciones: Revista de la Asociación Geológica Argentina, v. 25, p. 405–429. Azcuy, C.L., and di Pasquo, M.M., 1999, Carbonífero y Pérmico de las Sierras Subandinas, Cordillera Oriental y Puna, in Caminos, R., ed., Geología Argentina: Anales Instituto de Geología y Recursos Minerales 29, p. 239–260. Azcuy, C.L., Carrizo, H.A., and Caminos, R., 1999, Carbonífero y Pérmico de las Sierras Pampeanas, Famatina, Precordillera, Cordillera Frontal y Bloque de San Rafael, in Caminos, R., ed., Geología Argentina: Anales Instituto de Geología y Recursos Minerales 29, p. 261–317. Balistieri, P.R.M.N., 2003, Paleoicnologia da porção superior do Grupo Itararé na região de Mafra (SC): limitações paleoecológicas, paleoambientais e estratigráficas [Doctor of Science thesis]: São Leopoldo, Brazil, Universidade do Vale do Rio dos Sinos, 128 p.

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The Geological Society of America Special Paper 468 2010

Ichnology of the latest Carboniferous–earliest Permian transgression in the Paganzo Basin of western Argentina: The interplay of ecology, sea-level rise, and paleogeography during postglacial times in Gondwana Patricio R. Desjardins Luis A. Buatois M. Gabriela Mángano Department of Geological Sciences, University of Saskatchewan, Saskatoon, Saskatchewan, S7N 5E2, Canada Carlos O. Limarino CONICET (Consejo Nacional de Investigaciones Científicas y Técnicas) and Departamento de Ciencias Geológicas, Universidad de Buenos Aires, Pabellón 2, Ciudad Universitaria, C1428EHA, Buenos Aires, Argentina

ABSTRACT The late Paleozoic climatic evolution of Gondwana can be traced by analyzing the benthic ecology of its coastal environments as revealed by their ichnologic content. Latest Carboniferous–earliest Permian transgressive deposits occur in the lower member of the Tupe Formation, Paganzo Group, western Argentina. Three different trace-fossil assemblages are present as part of two complete depositional sequences. Trace-Fossil Assemblage 1 consists of Treptichnus pollardi and Helminthopsis abeli, which occur in fine-grained heterolithic facies. This assemblage characterizes a subaqueous freshwater substrate in a flood-plain environment. Trace-Fossil Assemblage 2, consisting of Halopoa isp., Palaeophycus crenulatus and Planolites montanus, occurs in thin-bedded, tabular sandstone. The tracemakers inhabited a low-energy distal-bay environment dominated by background sedimentation with sporadic storm episodes. The trace fossils represent the activity of post-storm colonizers. Trace-Fossil Assemblage 3 is monospecific, comprising only Rhizocorallium commune preserved at the interface between a sandstone bed and the overlying mudstone. The tracemakers inhabited the overlying muddy substrate in a low-energy distal-bay environment and burrowed down into the sediment, expanding laterally at the top of the underlying sandstone layer. Trace-Fossil Assemblages 2 and 3 do not resemble typical marine ichnocoenoses and can be considered a depauperate Cruziana ichnofacies, suggesting brackish-water conditions in a restricted marine embayment. The ichnofauna associated with the latest Carboniferous–earliest Permian transgression of the Tupe Formation is compared with that in the older (early Late Carboniferous) postglacial transgression recorded in the underlying Guandacol Formation. The latter reflects Desjardins, P.R., Buatois, L.A., Mángano, M.G., and Limarino, C.O., 2010, Ichnology of the latest Carboniferous–earliest Permian transgression in the Paganzo Basin of western Argentina: The interplay of ecology, sea-level rise, and paleogeography during postglacial times in Gondwana, in López-Gamundí, O.R., and Buatois, L.A., eds., Late Paleozoic Glacial Events and Postglacial Transgressions in Gondwana: Geological Society of America Special Paper 468, p. 175–192, doi: 10.1130/2010.2468(08). For permission to copy, contact [email protected]. ©2010 The Geological Society of America. All rights reserved.

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Desjardins et al. freshwater conditions related to an extreme meltwater influx coming from retreating glaciers in fjord environments. In contrast, the latest Carboniferous–earliest Permian transgression in the Paganzo Basin occurred in a confined, brackish-water embayment, but away from the direct influence of meltwater discharges.

INTRODUCTION Postglacial transgressions occur as a response to sea-level rise induced by the melting of ice masses. Thus, a salinity gradient is established within postglacial seas, controlled by the distance to freshwater discharge areas near to ice masses. However, salinity levels at a particular coastal environment can be also controlled by the local paleogeography. Salinity exerts a significant control on the distribution, abundance, and type of organisms in marginal-marine settings (Remane and Schlieper, 1971; McLusky, 1989). Accordingly, the ecology of the associated biota is highly influenced by freshwater discharges coming from the melting ice (Buatois et al., 2006, this volume). Trace fossils have proved to be a useful tool in determining paleosalinity levels (e.g., Pemberton and Wightman, 1992; MacEachern and Pemberton, 1994; Mángano and Buatois, 2004; Buatois et al., 2005; MacEachern and Gingras, 2007). During the late Paleozoic, Gondwana underwent three glacial and deglaciation events recorded in different basins of South America, South Africa, Antarctica, Australia, and India. These events had a major impact on coastal environments; deglaciation was accompanied by sea-level rises that triggered transgressive phases over Gondwana (López Gamundí, 1989; Limarino et al., 2002). The Paganzo Group of western Argentina in Cuesta de Huaco and adjacent areas provides an opportunity to analyze and compare the ichnology of two transgressive episodes. The first episode, recorded within the Guandacol Formation, is a transgression influenced by freshwater discharges from ice melting in a fjord environment (Limarino et al., 2002; Buatois and Mángano, 2003; Buatois et al., 2006). The second transgressive episode recorded, in the Tupe Formation, is contemporaneous with deglaciation phases in eastern basins in Gondwana, but does not show any direct influence or connection to melting ice masses. This paper analyzes the transgressive deposits within the lower member of the Tupe Formation and focuses on (1) ichnologic aspects of the latest Carboniferous–earliest Permian transgression; (2) environmental and paleogeographic controls affecting the ecology of this transgression, and (3) the comparison of the ichnofauna of the latest Carboniferous–earliest Permian transgression with that of the older, postglacial transgression (early Late Carboniferous) recorded in the Guandacol Formation in the same region.

(López Gamundí, 1997): (1) Late Devonian–Early Carboniferous (Solimões and Amazonas Basins in Brazil, and Lake Titicaca region of Bolivia); (2) early Late Carboniferous (Paganzo and Calingasta-Uspallata Basins of western Argentina and Tarija Basin of northwest Argentina and southern Bolivia); and (3) Late Carboniferous–Early Permian (various basins in Brazil, Uruguay, Paraguay, Antarctica, South Africa, India, and Australia). Each of these periods was followed by a postglacial sea-level rise, the establishment of fjords and lakes, and a subsequent climatic amelioration (e.g., Limarino and Césari, 1988; López Gamundí, 1989; Buatois and Mángano, 1994, 1995a; Limarino et al., 2002). The Paganzo Basin is a pericratonic foreland basin fringed on its western margin by the Protoprecordillera, a positive element that divided the drainage between the Calingasta-Uspallata and Paganzo Basins. However, marine deposits within the Paganzo Basin indicate the existence of a breach in the Protoprecordillera during the Late Carboniferous–Early Permian (López Gamundí et al., 1992). Our studied sections are located in the surroundings of Cuesta de Huaco, San Juan province (Fig. 1), where three formations are recognized in the Paganzo Group: Guandacol, Tupe, and Patquía (Fig. 2) (Limarino et al., 1986). The Guandacol Formation records a transition from a glacial to a postglacial period (Limarino et al., 2002; Pazos, 2002a). Coarse-grained diamictite formed in subaqueous outwash environments (Limarino et al., 2002; Marenssi et al., 2005) are replaced upward by mudstone and thinly laminated deposits containing dropstones. The upper interval of the Guandacol Formation mainly consists of sandstone and records the progradation of a Gilbert-type delta (Limarino et al., 2002). The Tupe Formation unconformably overlies the Guandacol Formation and consists of fluvial deposits punctuated by a short marine interval at its lower part (Cuerda, 1965; Limarino et al., 1986; Pazos, 1994; Desjardins et al., 2009). In particular, Desjardins et al. (2009) suggested the existence of a punctuated transgression, in which a transgressive coastal plain was interrupted by the establishment of a braided-fluvial system, which later evolved into a marginal-marine embayment. This marine interval has a limited distribution in the basin (Fig. 3), in contrast to the older (Namurian–early Westphalian) postglacial transgression. The Patquía Formation consists of conglomerate, sandstone, and evaporite recording fluvial, eolian, and playa-lake environments in an arid to semiarid setting (Limarino et al., 1986).

GEOLOGIC SETTING Like other late Paleozoic successions of Gondwana, the Paganzo Group records a complex paleoclimatic evolution, as a response to the movement of the supercontinent over high latitudes. Three glacial events have been identified in Gondwana

DEPOSITIONAL SETTING AND SEQUENCE STRATIGRAPHY OF THE STUDIED SECTIONS The lower interval of the Tupe Formation was studied in detail in the localities of Quebrada La Delfina and Mina La

Paganzo Basin of western Argentina

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Figure 1. Location of the studied outcrops (Mina La Ciénaga and Quebrada la Delfina) and simplified geology of the Cuesta de Huaco area.

Ciénaga (Figs. 1 and 4). Sixteen facies grouped in seven facies associations have been identified and described within two complete depositional sequences (Tables 1 and 2). The environmental framework of MacEachern and Gingras (2007) for bay successions is adopted in this study. These authors distinguished restricted bays that have limited or intermittent connection to the open sea, from open bays that have virtually unrestricted connection to the open sea. Transgressive deposits in the Tupe Formation are interpreted to record deposition in a restricted bay. The contact between the Guandacol Formation and the Tupe Formation corresponds to a sequence boundary (SB). Coastalplain deposits of the Tupe Formation unconformably overly

delta-front deposits included within the Guandacol Formation. A detailed sedimentologic and sequence-stratigraphic analysis of these deposits has been presented by Desjardins et al. (2009). Sequence 1 Lowstand Systems Tract Sheet-like, fining-upward sandstone packages, commonly displaying lateral-accretion surfaces (Facies Association I) are interpreted as high-sinuosity channels. Facies Association I also includes thinly interbedded sandstone, siltstone, mudstone, and coal facies, representing flood plains, peat, and associated water bodies.

Patquia Fm

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Transgressive Systems Tract Toward the top of the deposits included within Facies Association I, an increase in coal thickness from 1–4 cm to 1 m is observed. The thicker coal bed at the top indicates an increase in accommodation for peat production. The thicker coal beds may reflect a rise in the water table within a retreating shoreline. As in Mina La Ciénaga, coastal-lagoon deposits cap a coastal-plain interval, indicating that the upper parts of Facies Association I and Facies Association II are part of the transgressive systems tract. The transgressive surface is located at the base of the first thick coal bed included in Facies Association I (Fig. 4). The establishment of water bodies toward the top of this systems tract (Facies Association II) indicates that the water table rise exceeded the sediment influx and peat production (Bohacs and Suter, 1997). The maximum flooding surface is located within the coastallagoon facies of Mina La Ciénaga (Fig. 4). The trace fossils present in sequence 1 occur toward the top of the lowstand systems tract and the base of the transgressive systems tract. This ichnofauna includes Helminthopsis abeli and Treptichnus pollardi (Trace-Fossil Assemblage 1). Highstand Systems Tract Fine- to medium-grained, thickening-upward tabular sandstone packages interbedded with chloritic mudstone (Facies Association II) record progradation of clastic wedges into a water body. The chloritic composition indicates that the lagoon was periodically connected with the adjacent embayment, because chlorite dominates in marine intervals that have the Precordillera as their main source area (Net et al., 2002). Sandstone deposits are interpreted as river-fed quasi-steady turbidity currents. Association with small mountain rivers and the brackish-water nature of the lagoon promoted formation of hyperpycnal flows. This systems tract is bounded at the top by a sharp unconformity of regional extent and sequence-boundary hierarchy, which represents a change from coastal-plain to braided alluvial-plain settings (Fig. 4). Sequence 2

Fjord Embayment Subaqueous Outwash mudstone

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Wave ripple cross-lamination

Planar-lamination

Coal

Figure 2. General stratigraphy and sedimentary environments of the Paganzo Group in the western margin of the Paganzo Basin.

Lowstand Systems Tract Pebbly and coarse-grained to medium-grained amalgamated sandstone facies (Facies Association III), representing multistory channels and longitudinal bars, are preserved with variable thicknesses (0.4–2 m) in the studied localities. Lowstand deposits are relatively thin in Mina La Ciénaga compared to those in Quebrada La Delfina, indicating that sediment bypass took place at this locality. However, the Quebrada La Delfina lowstand deposits are up to 50 m thick, suggesting differential accommodation potential along depositional strike (Fig. 4). A 1-m-thick coal bed toward the top of this facies association indicates an increase in accommodation and water table (Fig. 4). Transgressive Systems Tract The lower portion of this systems tract shows different facies associations in the two localities. The transgressive surface in

Paganzo Basin of western Argentina Argentina

67° W

179

Fluvio-deltaic deposits

N

N

Shallow- marine and bay deposits

?

Shoreline at maximum transgression during the earliest Permian

ST EM

Las Gredas

SY

Cortaderas Malimán

Guandacol

Study area Huaco

La Rioja

S Sañ ierra oga sta

Rest

ricted

Open

emb

-marin

ayme

e

nt

FAM ATI NA

29° S

C H I L E

?

S I E R R A S PA M P E A N A S

RÍO BLA NCO BA SIN

Río del Peñon

PAGANZO BASIN Olta Malanzán

Sierra de Malimán

?

San Juan

SIERRA DE PIE DE PA L O

Figure 3. Paleogeographic map of the Paganzo Basin, and spatial distribution of the latest Carboniferous–earliest Permian transgression.

32° 30' S

0

Carboniferous Permian

Mendoza

100

Latest Carboniferous - earliest Permian transgression

Namurian postglacial transgression

Quebrada La Delfina is expressed by the presence of thick coal beds associated with other fine-grained flood-plain deposits, which are replaced upward by deltaic facies (Facies Association IV). In Mina La Ciénaga, the transgressive surface is of waveravinement nature, and bay-shoreline deposits overlie the braided

alluvial-plain deposits. The bay-margin facies are grouped together in parasequences 1, 2, and 3. Bay-margin deposits (Facies Association V) include a wide range of medium- to fine-grained sandstone facies (Sl, Swr, Sh, Sr [see Table 1 for definition of abbreviations]) and heterolithic facies. Interbedded mudstone and fine-grained sandstone facies (Facies Association VI), which

LST

SB P7

FA VI

Maximum flooding surface

TST

P3

FS FS

P2

FS

P1

FA III

Transgressive surface

FS

LST Maximum flooding surface

HST TST

Transgressive sur

face

FA III

FA II

Sequence boundary

FA I

Sequence 1

TFA 2

FA IV

Sequence 2

FS FS

P5 P4

TFA 2

FA V

TFA 2

P6

Sequence 2

HST

TFA 3

FS FA VI

FA VII

Sequence boundary

FA VII

Quebrada La Delfina

6 km

FA V

Mina La Ciénaga

TFA 1

Sequence Mudstone vf f m c vc Sandstone

cg

LST boundary

Current ripple cross-lamination

Gastropods

Trough crossstratification

Bivalves

Planar crossstratification

Palaeophycus crenulatus; Planolites montanus

Parallel lamination

Treptichnus pollardi; Helminthopsis abeli

Wave ripple cross-lamination

Halopoa isp.

Low-angle cross-stratification

Rhizocorallium commune

Coal

TFA 1

FA I

Brachiopods

Sequence 1

Legend

10 m

Mudstone vf f m c vc Sandstone

cg

Figure 4. Correlation between the two studied sections, and distribution of facies associations, sequence stratigraphic surfaces, parasequences, systems tracts, and trace-fossil assemblages within Sequences 1 and 2. FA—facies association; TFA—trace-fossil assemblage; P—parasequence; LST—lowstand systems tract; TST—transgressive systems tract; HST—highstand systems tract; SB—sequence boundary; FS—flooding surface. The panel is oriented roughly parallel to the coastline.

Climbing ripple cross-laminated, very well-sorted, fine- to very finegrained sandstone. Gradational and sharp bases and tops. Locally interbedded within thin layers of siltstone and claystone.

Planar to low-angle cross-stratified, well-sorted, medium- to finegrained sandstone. Sharp bases and tops. Parting lineation is present.

Wave ripple cross-laminated and low-angle cross-laminated, very wellsorted, very fine to fine-grained sandstone. Sharp bases and tops.

Very well-sorted, normally graded, fine- to very fine-grained sandstone. Sharp erosive base and dome-shaped tops.

Massive, well-sorted, medium- to very fine-grained sandstone. Gradational and sharp bases and tops.

Thin horizontally laminated mudstones, siltstone, and very fine-grained sandstone. Sharp bases and tops.

Thin horizontally laminated to massive rooted mudstone and siltstone. Sharp bases and tops.

Thinly laminated chloritic mudstone and siltstone. Sharp bases and tops. Massive and thinly laminated marine mudstone. Sharp bases and tops.

Sr

Sh

Swr

Srd

Sm

Fl

Fr

Fc

C

Coal and carbonaceous mudstone. Sharp bases and tops.

Sediment deposit and/or reworked by oscillatory currents

Horizontally laminated, well-sorted, very fine to medium-grained sandstone. Sharp to gradational bases and tops. Locally interbedded with thin layers of siltstone and claystone.

Sl

Fm

Sediment reworking by waves in the surf zone under an upper-flow regime

Planar cross-stratified, moderate- to well-sorted, very fine to mediumgrained sandstone. Erosive and sharp bases; sharp tops.

Sp

Organic material accumulation in setting with very low clastic sediment input

Suspension fallout

Suspension fallout

Coastal plain

Distal bay

Coastal lagoon

Coastal plain, alluvial plain

Soil development on alluvial and coastal plain deposits

Facies association: I, III

Facies association: V, VI

Facies association: II

Facies association: I, VII

Facies association: III, IV, VII

Coastal plain, alluvial plain, delta Suspension fallout, in times alternating with traction + suspension fallout from turbulent currents

Facies association: VI

Facies association: I, II, IV

Distal bay

Facies association: V, VI

Facies association: V

Facies association: I, II, III, IV, VII

Facies association: I, II, III, IV, VII

Facies association: I, III, IV, VII

Facies association: I, III, IV, V, VII

Facies association: III, VII

Facies association: III, VII

Facies association: III, VII

Distribution

Coastal plain, delta, coastal lagoon

Continuous traction + suspension fallout deposition from turbulent currents

Traction + suspension fallout from turbidity currents and high-energy waves

Bay margin and distal bay

Bay shoreline

Coastal plain, coastal lagoon, and delta

Braided river, coastal plain, coastal lagoon, and lake

Bed-load + suspension-load deposition from unidirectional currents Continuous traction + suspension fallout deposition from turbulent currents

Braided river, coastal plain

Tractive bed-load deposition from unidirectional currents in channels

Braided river, coastal plain, and bay margin

Tractive bed-load deposition from unidirectional currents

Trough cross-stratified, moderate- to well-sorted, medium- to coarsegrained sandstone. Sharp and erosive bases; sharp or gradational tops.

St

Braided river

Deposition from a concentrated density flow

Matrix-supported conglomerate. Well-rounded clasts scattered in a medium- to coarse-grained sandstone.

Gh

Braided and anastomosing rivers

Sedimentary environments Braided and anastomosing rivers

Tractive bed-load deposition from unidirectional currents as longitudinal bars and/or channel fills

Planar cross-stratified, gravelly very coarse-grained sandstone. Each set is limited by a sharp and planar base and top. Clay chips are common.

TABLE 1. FACIES CHART (BASED ON DESJARDINS ET AL., 2009) Lithology and sedimentary structures Sedimentary processes and depositional conditions Trough cross-stratified, gravelly very coarse-grained sandstone. Each Tractive bed-load deposition from set generally comprises a sharp and erosive base. Clay chips are unidirectional currents as channel fills common.

Gp

Gt

Lithofacies

Thick mudstone layers -Tabular, normally graded, and wave-rippled sandstone packages -Scarce marine fauna of the Tivertonia jachalensis– Streptorhynchus inaequiornatus Biozone -Ichnofauna: Rhizocollarium commune, Halopoa isp., Paleophycus crenulatus, Planolites montanus Erosionally based, laterally discontinuous sandstone bodies -Abundance of large (up to 1 m wide) clay chips -Thick siltstone beds

Sl, Sr, Sp, Sm, St, Fl

Sh, Swr, St, Fm Fm, Swr, Srd

Gh, Gt, Gp, Sp, St, Sr, Sl, Sm, Fl, Fr

IV: Interbedded medium- to fine-grained sandstone-mudstone lithofacies

V: Medium- to very fine-grained sandstone and heterolithic lithofacies

VI: Interbedded mudstone and finegrained bioturbated sandstone lithofacies

VII: Interbedded pebbly and coarsegrained sandstone-siltstone lithofacies

Distal bay

Anastomosing river

LST?

Bay shoreline and bay margin

River-dominated delta

Alluvial braided plain

Coastal lagoon associated to riverfed quasi-steady turbidity currents

TST to HST

TST

TST

LST to TST

TST to HST

Mina La Ciénaga Quebrada La Delfina

Mina La Ciénaga Quebrada La Delfina

Mina La Ciénaga Quebrada La Delfina

Quebrada La Delfina

Mina La Ciénaga Quebrada la Delfina

Mina La Ciénaga

Note: Based on Desjardins et al. (2009). LST—lowstand systems tract; HST—highstand systems tract; TST—transgressive systems tract. Lithofacies abbreviations are defined in Table 1.

Retrogradational stacking pattern of shallow-marine deposits -Dominance of wave-induced features

Laminated mudstone containing small plant fragments -Tabular fine-grained sandstone packages -Amalgamated sandstone bodies containing numerous reactivation surfaces

Laterally continuous, amalgamated lenticular package -Sharp-based, normally graded sandstone packages scouring into relatively thick coal lithofacies

St, Sp, Sr, Sl, Gt, Gp, Gh, Fl, C

III: Pebbly and coarse to medium-grained amalgamated sandstone lithofacies and coal lithofacies

Chlorite mineralogy of mudstone suggests a marine connection -Thickening-upward sandstone lithofacies (fluvial progradation)

Sl, Sm, Sr, Fc

II: Interbedded chlorite mudstone and sandstone lithofacies

TABLE 2. CHARACTERISTICS OF FACIES ASSEMBLAGES, LOCATION WITHIN THE SEQUENCE-STRATIGRAPHIC FRAMEWORK, AND ENVIRONMENTAL SIGNIFICANCE Facies associations Lithofacies Remarks Systems tracts Sedimentary Locality environment I: St, Sp, Sr, Fining-upward cycles LST to TST Coastal plain Mina La Ciénaga Interbedded medium- to fine-grained Sm, Sl, Fl, Fr, -Extensive, periodically inundated flood-plain Quebrada La Delfina sandstone-siltstone-coal lithofacies C deposits -Waterlogged soils -Thickening-upward coal -Ichnofauna: Helminthopsis abeli, Treptichnus pollardi

Paganzo Basin of western Argentina are interpreted as distal-bay deposits, make up parasequences 4, 5, 6, and 7. The tabular sandstone facies are interpreted as the product of a high-density flow, which probably originated either during a flooding stage of a river that flowed into the bay or from a storm-induced gravity flow. A maximum flooding surface is located at a mudstone horizon containing a fragmentary and scarce marine fauna of the Tivertonia jachalensis–Streptorhynchus inaequiornatus Biozone (Cisterna et al., 2005, 2006). In Quebrada La Delfina, brachiopods, bivalves, gastropods, and a few crinoids occur within this horizon. In Mina La Ciénaga, the only body fossils observed are present in a coquina composed of large gastropods. Within the distal-bay tempestites, four ichnotaxa were identified and grouped into two different assemblages. Halopoa isp., Palaeophycus crenulatus, and Planolites montanus are included in Trace-Fossil Assemblage 2, and Rhizocorallium commune compose TraceFossil Assemblage 3. These trace fossils occur only in a sandstone facies (Srd [see Table 1 for definition of abbreviation]). Highstand Systems Tract Poorly represented shallowing-upward distal-bay deposits occur above the maximum flooding surface. As a consequence of a base-level drop, a sequence boundary is located at the top of the distal-bay deposits, revealing the passage to continental environments in a semiarid climate (Facies Association VII). COMPOSITION AND CHARACTERISTICS OF THE ICHNOFAUNA Three different trace-fossil assemblages were identified in the stratigraphic sections. Each of these assemblages is composed of

183

all trace fossils occurring within a particular facies or association of facies. The integration of ichnologic and sedimentologic data provides the opportunity for refining interpretations of sedimentary environments and depositional evolution within a sequencestratigraphic framework. Trace-Fossil Assemblage 1 Trace-Fossil Assemblage 1 consists of Helmintopsis abeli and Treptichnus pollardi in flood-plain deposits. The assemblage is present at the base and top of very fine-grained sandstone (Sr) within the fine-grained heterolithic facies (Facies Association I). Four slabs containing different specimens were collected and analyzed. Within this facies, paleosols with distinctive rhizoliths are present. . Helminthopsis abeli Ksia˛zkiewicz, 1977 (Fig. 5A): Simple, irregularly meandering and winding horizontal trails. The meanders are irregular in shape and size. Diameter is 0.3–0.4 mm. The trace fill is similar to the host rock. All specimens are preserved as positive epirelief in mudstone facies. Helminthopsis is interpreted as the product of the combined activity of locomotion and feeding (grazing traces) by arthropods and nematodes (Buatois and Mángano, 1993a). Treptichnus pollardi Buatois and Mángano, 1993b (Fig. 5B): Simple burrow systems consisting of straight, unlined, unornamented, horizontal burrow segments. Small pits occur mostly within burrow segments and locally at the angle of juncture, representing the bedding-plane expression of vertical shafts. Diameter is 0.4–0.5 mm. Length of individual segments is 1.2– 5.3 mm. Pit diameter is 1.1–1.6 mm. Specimens preserved as positive and negative epirelief in mudstone facies. Treptichnus is

Figure 5. Trace-Fossil Assemblage 1. (A) Helminthopsis abeli. (B) Treptichnus pollardi.

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a feeding trace that in freshwater environments may be produced by worms or insect larvae (Buatois and Mángano, 1993b; Buatois et al., 1998a). The environmental implications of this trace-fossil assemblage are consistent with our interpretation for the associated facies. The tracemakers inhabited a subaqueous freshwater substrate in a flood-plain environment characterized by good oxygenation, and low-energy and sedimentation rates, periodically disturbed by overbank flooding. No evidence of marine connection has been detected for this facies assemblage. Other ichnofaunas from identical environmental settings of the Tupe Formation show similarities with the assemblage observed in the studied localities. Buatois and Mángano (2002) reported the occurrence of poorly specialized grazing and feeding trails (Helminthoidichnites tenuis, Planolites isp.) and other ichnotaxa, such as Archaeonassa fossulata, Didymaulichnus lyelli, and Palaeophycus tubularis, in deposits of the Tupe Formation in the vicinity of Huerta de Huachi. As in the case of Trace-Fossil Assemblage 1, the ichnofauna documented by Buatois and Mángano (2002) is characterized by the dominance of very simple forms, superficial or very shallow trace fossils, mostly restricted to bedding-plane surfaces, allowing preservation of the primary sedimentary fabric. Dominance of bedding-plane trace fossils and little disturbance of primary fabric are typical of Paleozoic flood-plain deposits (Buatois et al., 1998b). Flood-plain water-body trace-fossil assemblages show an expected low ichnodiversity in comparison with their equivalents from lacustrine basins. The depauperate assemblages in flood-plain deposits have been linked to the less stable conditions and the temporary nature of their ponds (Buatois and Mángano, 2002). Trace-Fossil Assemblage 2 Trace-Fossil Assemblage 2 consists of Palaeophycus crenulatus, Planolites montanus, and Halopoa isp. This assemblage occurs in relatively thin and tabular distal-bay sandstone packages (Sr) of Facies Association VI. Palaeophycus crenulatus Buckman, 1995 (Figs. 6A and 6B): Straight to sinuous, horizontal, unbranched, thinly lined, cylindrical, smooth-walled, endichnial burrows having distinctive annulations. Overcrossing is very common (Fig. 6A). Burrow diameter is 4.8–6.2 mm. Burrow fill matches the host rock and is massive. Annulations have a spacing of 0.9–1.6 mm (Fig. 6B). Pemberton and Frey (1982) and Mángano et al. (2002) reviewed Palaeophycus. Palaeophycus crenulatus is characterized by a lined open burrow with millimeter-scale annulations. The annulations indicate that the burrow lining was constructed as a series of ring-shaped structures. Palaeophycus is interpreted as the dwelling structure of predaceous or suspension-feeding animals (Pemberton and Frey, 1982). The polychaete Glycera has been proposed as a modern analogue for the Palaeophycus tracemaker by Osgood (1970). Planolites montanus Richter 1937 (Fig. 6C): Horizontal to subhorizontal, subcylindrical sinuous and undulating, unlined

trace fossils. Diameter is 2.4–2.7 mm and constant within a specimen. The trace fill differs from the host rock in grain size and color. The trace surfaces are typically smooth. Specimens are preserved as full relief. The ichnotaxonomy of Planolites has been addressed by Pemberton and Frey (1982), who noted that the diagnostic characteristic of P. montanus is the curved to contorted shape of the traces. These authors interpreted Planolites as feeding structures of deposits feeders, most likely polychaetes. Other authors (e.g., Fillion and Pickerill, 1990; Knaust, 2007), however, considered that other phyla could also be involved. In particular, Knaust attributed P. montanus to the activity of bivalves and other mollusks. Halopoa isp. (Figs. 6D–6F): Horizontal to oblique burrows comprising inflated-shape segments covered with longitudinal irregular striations. The segments are unbranched. Burrow segments are 10.8–21.1 mm wide and 8.9–22.8 mm long. Striae are 0.8–0.9 mm. Burrows are preserved as full relief on both tops and bases of sandstone beds. Halopoa has been revised by Uchman (1998, 2001), who recognized three different ichno. species: H. imbricata Torell, 1870; H. annulata Ksia˛zkiewicz, 1977; and H. storeana Uchman, 2001. Halopoa imbricata is unbranched, and has long and continuous furrows and wrinkles. Jensen (1997), however, placed Halopoa imbricata in Palaeophycus as P. imbricatus. Halopoa annulata is branched and contains perpendicular constrictions. Halopoa storeana has wrinkles arranged in a distinct, oriented plait-like design. The studied specimens display diagnostic features of Halopoa (e.g., predominately horizontal structures, inflated morphology, and longitudinal tension-related striations). However, determination at ichnospecific level is not possible. The fragmentary nature of the analyzed specimens prevents evaluations of the overall architecture. Two possible tracemakers may have produced Halopoa: (1) worm-like organisms that could expand its body hydraulically when moving through the sediment, and (2) crustaceans that could push against the walls with their carapace (Seilacher, 2007). The trace-fossil assemblage described is characterized by low ichnodiversity but high abundance. The degree of bioturbation is locally intense, mostly due to the presence of high-density assemblages of Halopoa isp. The assemblage consists dominantly of horizontal trace fossils produced by deposit and suspension feeders. The tracemakers presumably inhabited a low-energy distal-bay environment dominated by background sedimentation punctuated by storm episodes. However, the trace fossils are preserved in thin-bedded tempestites, most likely representing the activity of post-storm colonizers. Although Palaeophycus and Planolites are facies-crossing forms that occur in both continental and marine settings, Halopoa seems to be restricted to marine conditions (e.g., Seilacher, 1955; Crimes and McCall, 1995; . Ksia˛zkiewicz, 1977; Jensen, 1997; Uchman, 1998; Buatois et al., 2001; Mángano et al., 2002). The more intense bioturbation in the Halopoa beds is also a significant difference with respect to Trace-Fossil Assemblage 1; freshwater ichnofaunas show no disturbance of the primary sedimentary fabric.

Figure 6. Trace-Fossil Assemblage 2. (A) Palaeophycus crenulatus. (B) Close-up of one specimen showing distinctive annulations. (C) Planolites montanus. (D) Halopoa isp. (E) Halopoa isp. (F) Close-up of the striae of Halopoa isp.

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Trace-Fossil Assemblage 3 This is a monospecific trace-fossil assemblage that consists of Rhizocorallium commune. It occurs at the top of one bed of distal-bay tempestites (Facies Association VI) that can be traced laterally for tens of meters. Rhizocorallium commune Schmid, 1876 (Figs. 7A–D): Horizontal to subhorizontal, straight, relatively short U-shaped burrows displaying a protrusive spreite. The tube diameter is 3.0–9.1 mm, burrow width is 31.0–107.2 mm, and length is 90.1–240.2 mm (terminology following that of Schlirf, 2000). The spreite is continuous, having the same material as the host rock and abundant fecal pellets. The arms of the causative burrow may be parallel to each other or widen downward in some specimens. The pellets are cylindrical in shape and are 0.9–1.5 mm long, and 0.4–0.8 mm wide. The ichnogenus Rhizocorallium has been reviewed by Fürsich (1974). According to this author, three different ichnospecies can be recognized: R. jenense Zenker, 1836, R. irregulare Mayer, 1954, and R. uliarense Firtion, 1958.

Recently, Knaust (2007) reexamined the type area and distinguished R. commune in addition to the other three ichnospecies. Rhizocorallium jenense is distinguished from the other ichnospecies by consisting of a straight, short, U-shaped spreiten-burrow commonly oblique to the bedding plane. However, the burrow as a whole can be locally vertically retrusive. Rhizocorallium irregulare is characterized by long sinuous, bifurcating or planispiral U-shaped spreiten-burrows, whereas R. uliarense presents a diagnostic trochospiral U-shaped spreiten-burrow. Rhizocorallium commune is mainly horizontally oriented and exhibits a clear spreite. Crustaceans have been assigned as the most likely producers of Rhizocorallium by several authors (Weigelt, 1929; Seilacher, 1955, 1967; Fürsich, 1974; Schlirf, 2000). This idea is supported by a relatively wide causative tube, the common preservation of scratch marks (Fürsich, 1974), and the presence of cylindrical or rod-shaped pellets, which are commonly associated with the activity of crustaceans (Ekdale et al., 1984; Bromley, 1996). However, the wide range of morphologies and their occurrences

Figure 7. Trace-Fossil Assemblage 3. Rhizocorallium commune. (A–C) General view. Coin in C is 2.5 cm. (D) Close-up showing rod-like pellets organized in a spreiten structure.

Paganzo Basin of western Argentina in different paleoenvironments suggested either a combination of different deposit-feeding crustaceans or a mixture of crustaceans and worm-like animals (Knaust, 2007). Although the crustacean origin is the more widely accepted, it is still debated whether Rhizocorallium represents the activity of suspension- or depositfeeding organisms (Fürsich, 1974; Knaust, 2007; Seilacher, 2007). Specimens from the Tupe Formation are characterized by relatively short but dominantly horizontal burrows, and the general architecture closely resembles that of R. commune. Density of specimens is moderate over a single bedding plane (Fig. 7D). A wide range of sizes is observed, suggesting that the surface was available for colonization for a considerable amount of time. Specimens do not crosscut each other. Some large specimens (Figs. 7A and 7D) display downward expansion, suggesting growth of the producer. From the horizontal orientation, abundance of pellets, and the tightly constructed spreite, a depositfeeding mode of life is inferred. Trace-Fossil Assemblage 3 represents the activity of a benthic community that inhabited the muddy substrate of a lowenergy environment, burrowed down into the sediment, and expanded laterally at the top of the underlying sandstone layer. Specimens most likely penetrated into the fine-grained substrate from the overlying flooding surface. MacEachern and Gingras (2007) indicated that distal-bay deposits accumulate in the deepest parts of the bays or the most sheltered areas and tend to be mud-dominated, reflecting overall low-energy conditions. Rhizocorallium may be regarded as a facies-crossing ichnotaxon, although it is more common in shallow- to marginal-marine (e.g., Seilacher, 1955; Farrow, 1966; Hakes, 1976; El-Asa’ad, 1987; Fraaye and Werver, 1990; Mángano et al., 2002) than in deepmarine settings (e.g., Uchman, 1991). The only well-documented occurrence of Rhizocorallium in continental environments corresponds to firm-grounds and specimens are very different from those analyzed here (Fürsich and Mayr, 1981).

A

187

DISCUSSION Trace fossil have largely been used to infer salinity conditions in paleoenvironmental reconstructions (e.g., Wightman et al., 1987; Pemberton and Wightman, 1992; MacEachern and Pemberton, 1994; Buatois et al., 1997, 2005; Mángano and Buatois, 2004; MacEachern and Gingras, 2007). In the specific case of late Paleozoic Gondwanic successions, a number of studies have dealt with the relationships between ichnofauna and paleosalinity in South American basins (e.g., Buatois and Mángano, 1992, 1993a, 1995a, 1995b, 2002, 2003; Nogueira and Netto, 2001; Pazos, 2000, 2002b; Balistieri et al., 2002, 2003; Buatois and del Papa, 2003; Mángano et al., 2003; Buatois et al., 2006, this volume). Buatois et al. (2006, this volume) described and compared the ichnology of postglacial transgressions in different basins of Gondwana. They concluded that postglacial transgressions commonly contain freshwater ichnofaunas of the Mermia and Scoyenia ichnofacies that developed as a response to extreme freshwater discharges caused by the melting of ice masses (see also Buatois and Mángano, 2003). In the case of the early Late Carboniferous transgression in the Paganzo Basin recorded by the Guandacol Formation and coeval units, they noted the presence of a freshwater assemblage, characterized by nonspecialized grazing traces, arthropod trackways, and fish trails (Figs. 8 and 9A). Paleoenvironmental reconstructions for these deposits suggest that deglaciation led to the establishment of large freshwater bodies to the east and a marine incursion from the west that flooded valleys, creating a series of fjords along the coast (Limarino et al., 2002; Marenssi et al., 2005; Buatois et al., 2006, this volume). The presence of acritarchs in some beds, mostly in the western region, supports a marine connection with intermittent periods of brackish-water conditions. However, acritarchs do not occur in beds containing freshwater ichnofaunas, and moreover, trace fossils are associated with terrestrially derived

B

C

Figure 8. Freshwater trace-fossil assemblage present in the deposits of the early Late Carboniferous postglacial transgression. (A) Mermia carickensis. (B) Maculichna carboniferous. (C) Orchesteropus atavus.

A.

5 7 3

6

4 2

1

1. Helminthopsis tenuis 2. Undichna britannica 3. Helminthoidichnites tenuis 4. Mermia carickensis 5. Treptichnus bifurcus 6. Circulichnis montanus 7. Maculichna carboniferus

B.

Trace-Fossil Assemblage 1 Trace-Fossil Assemblage 2 2 1

Trace-Fossil Assemblage 3

1

3

2

1

1. Helminthopsis abeli 2. Treptichnus pollardi

1. Planolites montanus 2. Halopoa isp. 3. Palaeophycus crenulatus

1. Rhizocorallium commune

Figure 9. Postglacial transgressive scenarios. (A) Fjord environment influenced by extreme freshwater discharge from retreating glaciers. A freshwater ichnofauna occurs in the transgressive deposits (Guandacol Formation). (B) Coastal environment without direct influence of freshwater influx from melting ice masses. A freshwater ichnofauna is present in coastal lakes and temporally inundated flood plains (Trace-Fossil Assemblage 1). A brackish-water ichnofauna (TraceFossil Assemblages 2 and 3) occurs in distal-bay facies (Tupe Formation).

Paganzo Basin of western Argentina palynomorphs (Buatois and Mángano, 2003; Buatois et al., 2006, this volume). Grazing trails are also present in the Trace-Fossil Assemblage 1, but associated with flood-plain pond deposits toward the base of the Tupe Formation, recording the activity of a freshwater biota. The described Trace-Fossil Assemblages 2 and 3, present in the latest Carboniferous–earliest Permian transgressive deposits of the Tupe Formation, provide further evidence of paleosalinity during sea-level rises. The presence of Halopoa and Rhizocorallium clearly indicates marine conditions, and represents a significant departure with respect to the early Late Carboniferous postglacial transgression in the area (Fig. 9B). This is consistent with the presence of the marine fauna of the Tivertonia jachalensis–Streptorhynchus inaequiornatus Biozone. The discrepancy between the ichnofaunas of these two coastal settings is mainly due to differences in the salinity levels of their transgressive seas. The distance of the depositional area with respect to the melting ice masses that triggered the transgressions created a salinity gradient. The Guandacol Formation contains an ichnofauna dominated by simple trails and arthropod trackways, which developed as a consequence of the high freshwater influx coming from the melting ice. The Tupe Formation bay deposits recorded a transgression that occurred far away from any glacial center. If the coastal areas transgressed are far from any glacial center, a marine signature is clearly expressed in the associated ichnofauna. The transgressive deposits recorded in the Tupe Formation lack any evidence of glacial influence (e.g., dropstones), supporting previous schemes of paleoclimatic evolution that suggest climatic amelioration in the Paganzo Basin during the Carboniferous–Permian transition (López Gamundí et al., 1992). However, coeval transgressive deposits in the Paraná Basin display clear glacial signatures and suggest that the latest Carboniferous–earliest Permian transgression in Argentina is a glacioeustatic response to deglaciation in the eastern basins (Buatois et al., 2002). Several ichnologic features support the notion that the latest Carboniferous–earliest Permian transgression was not characterized by fully marine conditions in the study area. Overall, trace fossils are exceedingly rare throughout the section studied. Although the scarcity of bedding planes may be detrimental to trace-fossil identification, examination of vertical sections reveals little or no bioturbation but preservation of primary sedimentary structures in most of the beds. The trace fossils do not resemble typical offshore ichnocoenoses, which are commonly characterized by the archetypal Cruziana ichnofacies (Pemberton et al., 1992). Offshore environments under fully marine salinity conditions are commonly highly bioturbated and contain a high-diversity ichnofauna. In contrast, ichnodiversity in the marine deposits of the Tupe Formation is very low. Interestingly, the notable exception is the Halopoa-bearing deposits. In addition, providing that bedding-plane views are available, trace-fossil suites consist of only one or two ichnotaxa. The tracefossil assemblages present in the marine intervals of the Tupe Formation represent a depauperate Cruziana ichnofacies, reflect-

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ing stressed environmental conditions. The horizon containing a marine fauna is interpreted to contain the maximum flooding surface, and suggests that normal or near-normal marine conditions in this area of the basin were only reached during times of maximum transgression. Deposits of the latest Carboniferous–earliest Permian transgression in less-restricted areas (e.g., Quebrada La Herradura) contain a more abundant and diverse fauna than in the localities analyzed in this article. Studies in late Paleozoic and Mesozoic marginal-marine environments suggest that brackish-water trace-fossil assemblages are characterized by (1) low ichnodiversity, (2) forms typically found in marine environments, (3) mixture of vertical and horizontal trace fossils from the Skolithos and Cruziana ichnofacies, (4) dominance of infaunal traces rather than epifaunal trails, (5) simple structures produced by trophic generalists, (6) variable abundances, (7) presence of monospecific associations, and (8) small sizes (Wightman et al., 1987; Pemberton and Wightman, 1992; MacEachern and Pemberton, 1994; Mángano and Buatois, 2004; Buatois et al., 2005). Some of these features are present in Trace-Fossil Assemblages 2 and 3. In particular, TraceFossil Assemblage 2 is characterized by a few ichnotaxa, and is dominated by simple and relatively small infaunal trace fossils. The bioturbation degree is variable, being locally intense where Halopoa isp. is present. Trace-Fossil Assemblage 3 is a monospecific assemblage. The overall simple architecture and restricted distribution of these structures also suggest stressed conditions. The envisaged brackish-water scenario is consistent with the presence of abundant microbial mat textures in coeval deposits identified farther west in the Uspallata region by Buatois et al. (2002). Formation of microbial mats is commonly inhibited in Phanerozoic open-marine environments due to disruption by infaunal organisms. Buatois et al. (2002) noted that some marine intervals in the coeval Santa Elena Formation lack bioturbation and contain a wide variety of structures indicative of microbial binding (e.g., Manchuriophycus-like cracks, load-casted ripples) that suggest anomalous conditions, probably reflecting dilution of seawater. The origin of brackish-water conditions during the latest Carboniferous–earliest Permian is attributed to the local paleogeography of the basin. The basin was fringed in all directions by topographic highs, including the Protoprecordillera to the west, which severely limited water circulation between the more marine settings of the west and the embayment on the east (Fig. 3). A breach on the Protoprecodillera served as a connection between the two basins, but at the same time this topographic high allowed the Paganzo embayment to remain brackish. The latest Carboniferous–earliest Permian transgressive deposits of the Tupe Formation cannot be traced along depositional dip for long distances (i.e., more than 100 km). The inherited paleogeography from the glacial times favored the development of embayments, and the patchy distribution of these deposits along the entire basin supports the notion of embayments rather that an open sea. The scarcity of storm deposits, particularly in the baymargin successions, is consistent with deposition in a restricted

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bay. The presence of the Protoprecordillera would have protected the embayment from high-energy open-marine waves. Finally, another peculiarity of this embayment is the absence of tidal signatures, which suggests a microtidal regime in the late Paleozoic seaway. In addition to basin restriction, late Paleozoic Gondwanic seas only locally may have reached fully marine salinity conditions because of the overall high discharges of meltwater during postglacial episodes. Comparisons with Quaternary analogues are not straightforward mainly because ichnologic studies in glacial to postglacial settings are scarce. Virtasalo et al. (2006) documented a pattern of vertical distribution of trace fossils in Holocene cores from the Baltic Sea that resembles that of the Paganzo Basin, albeit at a smaller scale. An increase in ichnodiversity was noted toward the end of the lacustrine phase in relation to a marine influence. Glacial-fed freshwater Ancylus Lake deposits are replaced by those from the brackish-water Littorina Sea. The trace-fossil assemblage (Palaeophycus assemblage) of the Ancylus Lake consists of Palaeophycus and Arenicolites, representing an extremely depauperate ichnofauna. With the establishment of brackish-water conditions, the Palaeophycus assemblage was replaced by an assemblage consisting of Planolites, Arenicolites, Lockeia, and Teichichnus (Planolites assemblage). However, ichnodiversity levels remain very low as a result of brackish-water conditions and restricted oxygen availability. CONCLUSIONS 1. Ichnofaunas from the latest Carboniferous–early Permian Tupe Formation reflect the shift from a nonmarine freshwater community (Trace-Fossil Assemblage 1) to a marine community under brackish-water conditions (Trace-Fossil Assemblages 2 and 3). 2. Helminthopsis abeli and Treptichnus pollardi (TraceFossil Assemblage 1) occur in subaqueous freshwater substrates of small water bodies on temporally inundated flood plains of a transgressive coastal plain. This assemblage represents a depauperate Mermia ichnofacies. The lower diversity of the flood-plain–pond ichnofaunas in comparison with their equivalents of the Mermia ichnofacies of lacustrine and fjord environments is linked to the less stable conditions and temporary nature of the floodplain water bodies. 3. Halopoa isp., Palaeophycus crenulatus, and Planolites montanus (Trace-Fossil Assemblage 2) characterize a low-energy distal-bay environment dominated by background sedimentation. Their preservation in thin-bedded tempestites reflects the activity of post-storm colonizers. Rhizocorallium commune (Trace-Fossil Assemblage 3) results from the activity of infaunal deposit feeders, which inhabited the muddy distal-bay environment during maximum transgression. The tracemakers burrowed down from the maximum flooding surface, and expanded laterally at the interface with the underlying tabular sand-

stone layer. These two assemblages represent a depauperate Cruziana ichnofacies, suggesting brackish-water conditions. 4. In a postglacial transgressive scenario, the proximity to retreating ice masses is reflected in the nature of the ichnofauna. The postglacial transgression recorded in the underlying Guandacol Formation developed in a fjord environment subject to extreme freshwater influxes, and the associated ichnofauna corresponds to the freshwater Mermia and Scoyenia ichnofacies. The latest Carboniferous–earliest Permian transgression of the Tupe Formation represents the eustatic response to deglaciation of distal ice masses covering eastern regions of Gondwana, but did not evolve under the direct influence of extreme meltwater discharges. The marine ichnofauna of the Tupe Formation represents the activity of an infaunal community in a brackish-water, restricted embayment. ACKNOWLEDGMENTS Financial support for this study was provided by the Argentinean Agency of Scientific Research, Natural Sciences and Engineering Research Council (NSERC) Discovery Grants awarded to Mángano and Buatois, as well as by University of Saskatchewan startup funds awarded to Buatois. We are grateful to Dirk Knaust and Joonas Virtasalo for providing useful reviews. Alfred Uchman is thanked for comments on ichnotaxonomic aspects, Gabriela Cisterna for providing information on the associated invertebrate fauna, and Ramiro Salvatierra for fieldwork assistance. Special thanks to Brian R. Pratt for his assistance in photographing the specimens, and his comments on the manuscript. REFERENCES CITED Balistieri, P.R.M.N., Netto, R.G., and Lavina, E.L.C., 2002, Ichnofauna from the Upper Carboniferous–Lower Permian rhythmites from Mafra, Santa Catarina State, Brazil: Ichnotaxonomy: Revista Brasileira de Paleontologia, v. 4, p. 13–26. Balistieri, P.R.M.N., Netto, R.G., and Lavina, E.L.C., 2003, Icnofauna de ritmitos do topo da Formação Mafra (Permo-Carbonífero da Bacia do Paraná) em Rio Negro, Estado do Paraná (PR), Brasil, in Buatois, L.A., and Mángano, M.G., eds., Icnología: Hacia una convergencia entre geología y biología: Publicación Especial de la Asociación Paleontológica Argentina, v. 9, p. 131–139. Bohacs, K., and Suter, J., 1997, Sequence stratigraphic distribution of coaly rocks: Fundamental controls and paralic examples: American Association of Petroleum Geologists Bulletin, v. 81, p. 1639–1697. Bromley, R.G., 1996, Trace Fossils: Biology, Taphonomy and Applications: London, Chapman and Hall, 361 p. Buatois, L.A., and Mángano, M.G., 1992, Abanicos sublacustres, abanicos submarinos o plataformas glacilacustres? Evidencias icnológicas para una interpretación paleoambiental del Carbonífero de la cuenca Paganzo: Ameghiniana, v. 2, p. 323–335. Buatois, L.A., and Mángano, M.G., 1993a, Trace fossils from a Carboniferous turbiditic lake: Implications for the recognition of additional nonmarine ichnofacies: Ichnos, v. 2, p, 237–258. Buatois, L.A., and Mángano, M.G., 1993b, The ichnotaxonomic status of Plangtichnus and Treptichnus: Ichnos, v. 2, p. 217–224, doi: 10.1080/ 10420949309380095.

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South America) and the Great Karoo–Kalahari Basin (Southern Africa): Ichnological implications: Gondwana Research, v. 5, p. 619–640, doi: 10.1016/S1342-937X(05)70634-1. Pemberton, S.G., and Frey, R.W., 1982, Trace fossil nomenclature and the Planolites-Palaeophycus dilemma: Journal of Paleontology, v. 56, p. 843–881. Pemberton, S.G., and Wightman, D.M., 1992, Ichnological characteristics of brackish water deposits, in Pemberton, S.G., ed., Applications of Ichnology to Petroleum Exploration: Society of Economic Paleontologists and Mineralogists, Core Workshop 17, p. 141–167. Pemberton, S.G., MacEachern, J.A., and Frey, R.W., 1992, Trace fossils facies models: Environmental and allostratigraphic significance, in Walker, R.G., and James, N.P., eds., Facies Models: Response to Sea Level Change: Waterloo, Geological Association of Canada, p. 47–72. Remane, A., and Schlieper, C., 1971, Biology of Brackish Water: New York, John Wiley and Sons, 372 p. Richter, R., 1937, Marken und spuren aus allen Zeiten, Parts 1 and 2: Senckenbergiana, v. 19, p. 150–169. Schlirf, M., 2000, Upper Jurassic trace fossils from the Boulonnais (northern France): Beringeria, v. 34, p. 145–213. Schmid, E.E., 1876, Der Muschelkalk des ostlichen Thuringen: Jena, Fromann, 20 p. Seilacher, A., 1955, Spuren und Fazies im Unterkambrium; in Beiträge zur Kenntnis des Kambriums in der Salt Range (Pakistan), in Schindewolf, O.H., and Seilacher, A., eds., Akademie der Wissenschaften und der Literatur zu Mainz, mathematisch-naturwissenschaftliche Klasse, Abhandlungen, v. 10, p. 11–143. Seilacher, A., 1967, Bathymetry of trace fossils: Marine Geology, v. 5, p. 413– 428, doi: 10.1016/0025-3227(67)90051-5. Seilacher, A., 2007, Trace Fossil Analysis: Berlin, Springer, 226 p. Torell, O., 1870, Petrificata Suecana Formationis Cambricae: Lunds Universitets Årsskrift 6: Afd, v. 28, p. 1–14. Uchman, A., 1991, Trace fossils from stress environments in CretaceousPalaeogene of Polish Outer Carpathians: Annales Societatis Geologorum Poloniae, v. 61, p. 207–220. Uchman, A., 1998, Taxonomy and ethology of flysch trace fossils: Revision of the Marian Ksiazkiewicz collection and studies of complementary material: Annales Societatis Geologorum Poloniae, v. 64, p. 105–208. Uchman, A., 2001, Eocene flysch trace fossils from the Hecho Group of the Pyrenees, northern Spain: Beringeria, v. 28, p. 3–41. Virtasalo, J.J., Kotilaine, A.T., and Gingras, M.K., 2006, Trace fossils as indicators of environmental change in Holocene sediments of the Archipelago Sea, northern Baltic Sea: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 240, p. 453–467, doi: 10.1016/j.palaeo.2006.02.010. Weigelt, J., 1929, Fossile Grabschächete brachyurer Decapoden als Lokalgeschiebe in Pommer und das Rhizocorallium-Problem: Z. f: Geschiebeforschung, v. 5, p. 1–42. Wightman, D.M., Pemberton, S.G., and Singh, C., 1987, Depositional modelling of the Upper Mannville (Lower Cretaceous), east-central Alberta: Implications for the recognition of brackish water deposits, in Tillman, R.W., and Weber, K.J., eds., Reservoir Sedimentology: Society of Economic Paleontologists and Mineralogists Special Publication 40, p. 189–220. Zenker, J.C., 1836, Historisch-topographisiches Taschenbuch von Jena and seiner Umgebung besonders in seiner naturwissenschaftlicher und medicinischer Beziehung: Jena, J. C. Zenker, 338 p.

MANUSCRIPT ACCEPTED BY THE SOCIETY 21 DECEMBER 2009

Printed in the USA

The Geological Society of America Special Paper 468 2010

Reconstruction of a high-latitude, postglacial lake: Mackellar Formation (Permian), Transantarctic Mountains Molly F. Miller Department of Earth and Environmental Sciences, Vanderbilt University, Nashville, Tennessee 37235, USA John L. Isbell Department of Geosciences, University of Wisconsin, Milwaukee, Wisconsin 53201, USA

ABSTRACT The Lower Permian Mackellar Formation is well exposed in a 10,000 km2 outcrop belt in the Nimrod, Beardmore, and Shackleton Glacier areas of the Transantarctic Mountains. This formation directly overlies glacial deposits and provides a unique glimpse of high paleolatitude conditions during the last icehouse to greenhouse transition. The unit records deposition in the Mackellar Lake or Inland Sea (MLIS), a fresh-water body at ~80° S paleolatitude that was broadly analogous to Glacial Lake Agassiz and was filled by fine-grained turbidites. Low total organic carbon (TOC) content and predominant vitrinite and inertinite are consistent with a low influx of organic matter from a sparsely vegetated, recently deglaciated terrain. A widespread but low-diversity ichnofauna and variable (although low overall) levels of bioturbation suggest oxic conditions and a bottom fauna restricted to areas of low sedimentation. Integration of sedimentologic, organic geochemical, and paleobiologic information with results of climate models and characteristics of modern lakes enhances reconstruction of parameters that controlled the functioning of the lake as an ecosystem. Regression equations relating mean annual temperature and mean depth of modern lakes to the number of ice-free days applied to the MLIS indicate ice cover from two to five months a year. Estimates of the depth of mixing and depth to the thermocline, based on maximum length, maximum width, and area, suggest a mixing depth of ~50 m and a thermocline of ~20 m. The MLIS probably was stratified during the summer and was dimictic, with overturns occurring after fall cooling and after ice melt; mixing was enhanced by turbidity currents. Productivity was low, as recorded by the low TOC, but organic matter fixed in the surface water of the lake may have been degraded and not recorded in the sediments. In spite of its high paleolatitude, the MLIS as reconstructed was dynamic and biologically active; the same probably was true of other Permian postglacial lakes.

Miller, M.F., and Isbell, J.L., 2010, Reconstruction of a high-latitude, postglacial lake: Mackellar Formation (Permian), Transantarctic Mountains, in LópezGamundí, O.R., and Buatois, L.A., eds., Late Paleozoic Glacial Events and Postglacial Transgressions in Gondwana: Geological Society of America Special Paper 468, p. 193–207, doi: 10.1130/2010.2468(09). For permission to copy, contact [email protected]. ©2010 The Geological Society of America. All rights reserved.

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INTRODUCTION Upper Paleozoic sedimentary rocks record the history of the last large-scale period of glaciation and subsequent warming, thus providing the most recent record of the transition from icehouse to greenhouse conditions (e.g., Rees et al., 2002; Powell, 2005). On a worldwide basis, this record is best exposed in postglacial lake deposits that occur on diverse Gondwana continents, including South America (Buatois and Mangano, 1995; López-Gamundí, 1997; Limarino et al., 2002;Trosdtorf et al., 2005; Archanjo et al., 2006), Africa (Visser, 1994; Scheffler et al., 2003), Australia (Dickins, 1996; Lindsay, 1997; Eyles et al., 2003; Jones and Fielding, 2004), India (Maejima et al., 2004), and Antarctica (Collinson et al., 1994; Miller and Collinson, 1994; Isbell et al., 2003). The well exposed Mackellar Formation of the central Transantarctic Mountains allows reconstruction of the postglacial events and history that characterized southern polar paleolatitudes immediately following the final melting of ice sheets in Antarctica during the earliest Permian. The Mackellar Formation was deposited in a large lake or series of smaller time-transgressive lakes that extended >1000 km through the present-day Transantarctic Mountains from north of the Nimrod Glacier to the Ellsworth Mountains (Fig. 1); the Mackellar body of water thus is broadly comparable to large Quaternary lakes of North America such as glacial Lake Agassiz. In this paper we integrate sedimentological, geochemical, and biological data to delineate the characteristics of the lake recorded by the Mackellar Formation and

Figure 1. Extent of outcrop of the Mackellar Formation and its correlatives from the Nimrod Glacier area to the Ellsworth Mountains.

compare them to features of modern lakes in order to reconstruct the ecosystem of the Makellar Lake or Inland Sea (MLIS). The focus herein is the outcrop between the Nimrod Glacier and the Shackleton Glacier (Fig. 2). During the Early Permian, this area lay at 75° S to 85° S, similar to its present location (Powell and Li, 1994). Permian to Jurassic sedimentary rocks in the central Transantarctic Mountains form a sequence several kilometers thick of continental glacial, lacustrine, fluvial, and coal-bearing deposits (Table 1). The Mackellar Formation consists of interbedded shale and sandstone and overlies the earliest Permian Pagoda Formation, which is composed of glacial deposits, primarly glaciolacustrine or glaciomarine, and secondarily subglacial deposits (Elliot, 1975; Barrett et al., 1986; Miller, 1989; Isbell et al., 2003). It is overlain by braided stream deposits of the Fairchild Formation, and coal-bearing fluvial deposits of the Buckley Formation. In situ stumps within the Buckley Formation reflect dense forests of Glossopteris trees that averaged >15 m tall (Knepprath et al., 2004), attesting to the climate amelioration that occurred after glaciation in spite of the continued high paleolatitude (Powell and Li, 1994). MACKELLAR FORMATION: DEPOSITIONAL SETTING AND PROCESSES The Mackellar Formation records the filling of a postglacial basin occupied by the Mackellar Lake or Inland Sea (MLIS). Its origin, setting, and size are comparable to those of glacial Lake Agassiz, although the near absence of shoreline features, paucity of internally correlative units, and poor age control preclude the detailed reconstruction that has elucidated the history of Lake Agassiz (e.g., Teller and Clayton, 1983; Teller and Bluemle, 1983; Teller, 2001; Teller et al., 2005). Excellent exposures of the Mackellar Formation in many areas of the Transantarctic Mountains allow interpretation of the depositional setting and processes, which are inferred to resemble those of Lake Agassiz. Major loci of deposition in Lake Agassiz were large fans of sediment at the mouths of spillways (Fenton et al., 1983). These fans record bed-load deposition from underflow currents (Fenton et al., 1983). The Mackellar Formation consists primarily of facies interpreted as underflow deposits that are analogous to underflow deposits of Lake Agassiz and that accumulated in diverse components of what has been described as a lacustrine turbidite sequence (Table 2; Fig. 3; Miller and Collinson, 1994). In the Beardmore Glacier area the unit is dominated by the interbedded sandstone and shale facies, with the associated shale facies and massive sandstone facies providing important information about the sedimentary processes. The shale facies records quiet-water conditions. In contrast, the interbedded sandstone and shale facies reflects quiet-water conditions interrupted by turbiditycurrent events that deposited upward-fining sandstones (Figs. 4 and 5; Tb-d of Lowe, 1982). In the Beardmore Glacier area, individual upward-fining beds are organized into upward-coarsening

Mackellar Formation, Transantarctic Mountains 170°W

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Figure 2. Outcrop locations within the study area. The shoreline of the Mackellar Lake or Inland Sea (MLIS) is inferred to have extended from north of the Nimrod Glacier away from the Ross Ice Shelf and south toward Mount Bowers.

TABLE 1. STRATIGRAPHIC UNITS AND DEPOSITIONAL ENVIRONMENTS, BEARDMORE GLACIER AREA Formation Rock type Thickness Inferred environment (fossil) (m) Buckley Formation Sandstone, shale, coal (Glossopteris) 750 Braided stream systems, lowland lakes Fairchild Formation Sandstone 250 Braided stream Mackellar Formation Shale, sandstone 100 Lake, inland sea 200 Glacial, glacial basinal Pagoda Formation Diamictite, sandstone Note: From Barrett et al. (1986), Isbell and Collinson (1988), Miller (1989), and Isbell et al. (2003).

sequences a few to tens of meters thick (Fig. 4). These packages are interpreted as crevasse splays of turbidite channels. The largest scale upward-coarsening sequences are capped by the massive sandstone facies. This facies is composed of beds up to 3 m thick of homogeneous, structureless fine-grained sandstone that fills cross-cutting channels and grades upward into sandstone with planar lamination and climbing ripple lamination (Fig. 6; Miller and Collinson, 1994). We infer that this sandstone facies was deposited in subaqueous channels that served as the major conduits of the turbidite system. At two localities the large-scale, cross-stratified sandstone facies records Gilbert-type deltas feeding the turbidite system, and the rare diamictite facies reflects

remobilization of poorly sorted glacial debris from the underlying Pagoda Formation. The burrowed sandstone facies records shallow shoreline deposition near the lake margin but away from centers of deposition. The shoreline of the MLIS is inferred to have swung east from north of the Nimrod Glacier and south along the present-day boundary between the Transantarctic Mountains and the ice landscape of the Polar Plateau (Fig. 2). In this configuration the Moore Mountains were closest to the locus of deposition, and the Shackleton Glacier area (Mount Butters, Ramsey Glacier) was least affected by sediment influx. The rate of deposition was very high in the Moore Mountains (Fig. 2), as evidenced by omnipresent

Sand: fine to very fine Mud: clay to coarse silt

Fine sand, with minor coarser component

Fine sand

Fine sand; minor clay, silt

Silt-sand matrix with clasts up to 3m

Interbedded sandstone and shale Abundant; dominant facies

Massive sandstone Present at all localities

Burrowed sandstone Rare

Large-scale crossstratified sandstone Rare

Diamictite Rare

Fills broad, shallow channel at one locality

Normally graded foresets with climbing ripple lamination

Bioturbation; symmetrical ripple marks

Massive, but grades up into planar and climbing ripple lamination. Fills broad channels, <5 m deep, hundreds of m wide

Normal grading, horizontal lamination, climbing ripple lamination, parting lineation, linguoid and 2-D ripples, local bioturbation

~3 m

10 m

10 m

Few meters

Tens of meters

Surrounded by interbedded sandstone and shale facies

Near top, grades into Fairchild Fm.

Near top, surrounded by interbedded sandstone and shale facies

Caps upward-coarsening sequences

Gradationally overlies shale. Includes hierarchy of upward-coarsening sequences, largest capped by massive sandstone facies

TABLE 2. MACKELLAR FORMATION FACIES AND FACIES INTERPRETATION Structures/geometry Scale Trends/association Horizontal lamination, regular to variable Few meters Most common at base of upwardthickness, normal grading coarsening sequences

Note: From Miller and Collinson (1994). 2-D—two-dimensional.

Grain size Clay, silt

Facies/abundance Shale Widespread, but minor

Deposited by debris flow locally derived from remnant highs of glacial diamictite

Possible Gilbert-type delta

Deposited above wave base, away from loci of deposition

Deposition from high and low (minor) turbidity currents

Deposition from high- and low-density turbidity currents

Interpretation Deposition from suspension and low-density turbidity currents

Figure 3. Depositional environment of the Mackellar Formation within the context of the underlying and overlying units (Table 1). Letters indicate inferred locations within the MLIS (Mackeller Lake or Inland Sea) of facies well exposed at outcrops shown in Figure 2. A—Moore Mountains; B—Mount Weeks; C— Tillite Glacier; D—Mount Butters and Ramsey Glacier; E—Mount Bowers.

A Figure 4. (A) Upward-fining sandstone in interbedded sandstone and shale facies, Tillite Glacier; 15 cm ruler for scale. (B) Upward-coarsening sequence in interbedded sandstone and shale facies, Tillite Glacier, recording progradation of crevasse splay of turbidite channel.

B

A Figure 5. (A) Interbedded sandstone and shale facies at Mount Butters, lacking well-defined upward coarsening; interpreted as deposited away from crevasse splay. (B) Fine sandstone with planar beds and climbing ripple lamination, indicating a high flow regime and rapid sedimentation near sediment source; Moore Mountains.

B

B A Figure 6. (A) Massive sandstone–filled channel (sandstone facies) that grades upward into climbing ripple lamination, Moore Mountains. Person in center for scale. (B) Permian marine turbidite, intensely bioturbated in both vertical and horizontal dimensions. Bluff, New Zealand. The penetrative and pervasive bioturbation contrasts with that of the Mackellar Formation (Figs. 7 and 8). Ruler is 1 cm long.

Mackellar Formation, Transantarctic Mountains climbing ripple lamination and the absence of bioturbation (Fig. 5B). In contrast, the Mackellar Formation at Mount Butters and the Ramsey Glacier in the Shackleton Glacier region consists of shale and interbedded sandstones that are not organized into upward-coarsening sequences (Fig. 5A). We interpret this region as the most distal and relatively unaffected by the crevasse splays of the turbidite system. Mount Bowers, near the Polar Plateau, is inferred to have been relatively close to the paleoshoreline; a basin-margin setting is consistent with paleocurrent data, suggesting transverse flow into the basin (Miller and Collinson, 1994, Fig. 2) and with inferred basin morphology (Isbell, 1999). BIOGENIC STRUCTURES AND DISTRIBUTION OF BIOTURBATION Trace fossils are relatively rare in the Mackellar Formation and consist primarily of simple horizontal endostratal trails assignable to the ichnogenera Paleophycus or Planolites (Fig. 7). Rarely the traces formed looped patterns on bedding planes, resembling the trace fossil Mermia (Buatois and Mangano, 1995); the distinctive traces Cochlichnus and Treptichnus also occur in rare horizons (Fig. 7). Scratch marks are discernible on the margins of some traces; some of these are bilobed and were produced by animals with skeletalized appendages (e.g., arthropods). Scratch marks occur on parts of otherwise structureless traces, implying preservational control over their distribution. The widespread, if uneven, distribution of diverse simple traces suggests that many, if not all, of the trace fossils were produced by arthropods. Some of the small traces may have been produced by conchostracans, which have been reported from proglacial lake deposits of the Pagoda Formation at Mount Butters (Fig. 2; Babcock et al., 2002).These conchostracans are known to have produced bilobed trails (Tasch, 1964). Bioturbation on bedding plane and vertical surfaces of the Mackellar Formation in the Shackleton and Beardmore Glacier areas has been assessed (Miller et al., 2002; Miller and Laban-

A

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deira, 2002; Miller et al., 2005) using pattern recognition methods (ichnofabric indicies [ii] of Droser and Bottjer, 1986; Bedding Plane Bioturbation Indices [BPBI] of Miller and Smail, 1997). Overall, there is relatively little bioturbation on horizontal surfaces of the interbedded sandstone and shale facies (Fig. 8), with 89% of observations showing no bioturbation (BPBI = 1) and only 6.3% of observations showing BPBI = 5 (60%–100% bioturbated). There is virtually no vertical bioturbation. Although the sparse bioturbation is widely distributed, important differences in the amount of bioturbation on bedding plane surfaces in different areas of the lake system (Fig. 8). In the Moore Mountains (Fig. 2) there is no bioturbation recorded on bedding plane surfaces. This area is closest to the lake margin, which was north of the present-day Nimrod Glacier, and there is abundant evidence of rapid sedimentation that probably precluded colonization by infaunal animals (Fig. 5B). Tillite Glacier, with well-developed upward coarsening sequences, records deposition close to turbidite channels in areas affected by crevasse splays (Figs. 2 and 4). Crevasse splay deposition and progradation are recorded by upward coarsening sequences and would be inimical to infaunal animals, resulting in the observed low levels of bioturbation. Extensive bioturbation would be most likely to have occurred in the distal parts of the lake, far from sources of sediment (e.g., in the Shackleton area), and also in nearshore areas that were protected from rapid sedimentation. In modern lakes, shallow littoral zones have the highest diversity of benthic animals, particularly in areas shielded from wave activity. The presence of symmetrical ripples at Mount Weeks indicates wave activity in a nearshore zone, but the relatively extensive bioturbation implies that this shoreline zone was sufficiently protected to be colonized by active burrowers. Levels of bioturbation at Mount Bowers are surprisingly high, with 21.6% of observations recorded as BPBI = 5. This may reflect longer periods between sedimentation events, as well as greater colonization by infaunal animals. Mount Bowers may have been close to the original shoreline (Fig. 2) and subject to enhanced deposition of organic

B

Figure 7. (A, B) Horizontal bioturbation in interbedded sandstone and shale facies, Mount Bowers. The bioturbation does not penetrate more than a few millimeters vertically. This degree of bioturbation is unusual for the Mackellar Formation. Ruler in mm.

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90 Mt. Bowers

80

Shackleton Gl Mt. Weeks Tillite Gl

% of observations

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Figure 8. Bioturbation on bedding planes at diverse outcrops of the Mackellar Formation within the study area. BPBI (Bedding Plane Bioturbation Indicies, Miller and Smail, 1997) are as follows: category 1, no bioturbation; BPBI 2, 0%–10%; BPBI 3, 10%–40%; BPBI 4, 40%; BPBI 5, 60%–100%. Number of observations: Mount Bowers, 1102; Shackleton Glacier area, 1188; Mount Weeks, 250; Tillite Glacier, 559; Moore Mountains, 1220. Gl—Glacier.

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matter. If so, the additional nutritional supply may have stimulated colonization by infaunal animals, resulting in more observations of intensely disrupted sediments (BPBI = 5; Fig. 8). The distribution of bioturbation indicates that bioturbation, a proxy for infaunal activity, is least in areas of high sedimentation, either near the shoreline or close to turbidity channels, and reaches highest levels farthest from loci of deposition toward the distal part of the lake (e.g., Mount Butters and Ramsey Glacier in the Shackleton Glacier area, Fig. 2). Bioturbation is also enhanced near the lake margins, in regions sheltered from sediment influx (e.g., Mount Butters and Mount Bowers) where organic matter may be less diluted by clastic sediment than elsewhere in the system, and where shallow depth and associated factors enhance benthic diversity (Kalff, 2002). The generally low level of bioturbation and its distribution mirrors that in modern lakes (White and Miller, 2008). PALEOSALINITY The Mackellar Formation in the Beardmore and Shackleton Glacier areas was deposited in fresh-water or slightly brackishwater conditions. This interpretation is based on three lines of evidence: (1) the presence of low-diversity ichnofauna of simple crawling traces, including some forms (e.g., Cochlichnus, Fig. 7B) that are known to be fresh-water ichnogenera and occur in lacustrine deposits of the (slightly) younger Buckley Formation; (2) high carbon/sulfur ratios of the shales, which are dis-

tinctly different from those of marine shales (Berner and Raiswell, 1984; Miller and Collinson, 1994); (3) the absence of marine trace fossils, including penetrative forms typical of marine turbidites (Fig. 6B); and (4) the absence of the Eurydesma fauna of marine body fossils that is characteristic of Permian high southern paleolatitude marine deposits (Banks and Clarke, 1987). The Mackellar Formation and correlative units can be traced >1000 km through the Transantarctic Mountains through the Ohio Range to the Ellsworth Mountains (Fig. 1). The Polarstar Formation in the Ellsworth Mountains, which is thicker than the Mackellar Formation and records sedimentation in a subsiding co-eval basin, contains marginal-marine trace fossils (e.g., Phycodes) but lacks marine body fossils, including the cold-water bivalve Eurydesma (Collinson et al., 1992). The Discovery Ridge Formation of the Ohio Range is also correlative with the Mackellar and Polarstar Formations (Collinson et al., 1994). The salinity conditions under which this formation was deposited are not clear, but marginal-marine trace fossils occur in the overlying unit, suggesting a similar origin for the Discovery Ridge Formation (Bradshaw et al., 1984; Aitchison et al., 1988). Salinity conditions inferred from these correlative units are consistent with a model of an inland sea from the (fresh-water) Mackellar Formation in the present-day Transantarctic Mountains to marginalmarine conditions in the present-day Ohio Range and Ellsworth Mountains. Although putative marine trace fossils indicate a marine influence in the Ohio Range and Ellsworth Mountains, the absence of the indisputably marine Eurydesma body fossil

Mackellar Formation, Transantarctic Mountains assemblage suggests that salinity conditions were not normal marine. The emerging scenario is one of an inland sea with a restricted opening to the paleo–Pacific Ocean in the Ellsworth Mountains area, becoming increasingly fresh-water–dominated down the (present-day) Transantarctic Mountains with fully fresh-water conditions in the present-day Shackleton and Beardmore Glacier areas. ORGANIC GEOCHEMISTRY Samples from the Mackellar Formation were analyzed for total organic carbon (TOC), type of organic matter, and vitrinite reflectance by DGSI using standard protocols (Miller et al., 1988); the results of TOC and vitrinite reflectance analyses are given in Table 3. Visually determined kerogen types from 10 samples indicate the presence of vitrinite (30%), inertinite (51%), and amorphous organic matter (19%) (Miller et al., 1988); similar percentages of kerogen types (30% vitrinite, 50% inertinite, 20% amorphous) were found by Horner and Krissek (1991, their fig. 2). Vitrinite is derived from terrestrial plant material (Hutton et al., 1994). Inertinite is derived from organic matter that was burned soon after deposition, extensively oxidized by bacterial degradation, or thermally altered. Distinguishing between these origins is possible, particularly if heating has not been too intense (e.g., Diessel, 1992; Mukhopadhyay, 1992), but this determination was not undertaken for the Mackellar Formation samples. In general, the Mackellar samples have little amorphous organic matter, suggesting deposition in an oxic or suboxic environment in which oxygen-consuming bacteria degraded organic matter. The widespread occurrence of biogenic structures implies oxygenated bottom conditions. The few samples high in amorphous organic matter are low in TOC. This is consistent with deposition under oxic conditions; under low-oxygen conditions, more organic matter would be preserved. The Mackellar Formation is intruded by mafic Jurassic sills. Horner and Krissek (1991, their figs. 15 and 19) demonstrated that vitrinite reflectance is raised and TOC lowered up to a distance equivalent to the sill thickness from the contact with a sill. This is consistent with studies suggesting that effects of intrusion on the organic matter diminish significantly in short distances from the contact (Clayton and Bostick, 1986). At Mount Bowers, vitrinite reflectance drops from 3.9 Ro to 1.71 Ro in <100 m stratigraphically; this is interpreted to reflect increasing distance from a Jurassic sill of unknown thickness presumed to be beneath the exposure (Miller et al., 1988).

TABLE 3. TOTAL ORGANIC CARBON (TOC) AND VITRINITE REFLECTANCE (Ro) OF SAMPLES FROM THE MACKELLAR FORMATION, BEARDMORE GLACIER AREA Number of Mean Standard Variance samples deviation Ro 12 2.88 0.73 0.53 TOC 20 0.61% 0.63 0.39

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The relationship between thermal alteration, as indicated by vitrinite reflectance, and TOC in the Mackellar Formation, is not clear (Miller et al., 1988; Horner and Krissek, 1991, their table 5). The mean Ro for 36 samples analyzed by Horner and Krissek (1991, their table 5) is 2.73, and the mean TOC of 107 samples is 0.322%. Samples from Mount Bowers (Fig. 2) have a much higher TOC than samples from other areas; all of the four Mount Bowers samples analyzed by Miller et al. (1988) have TOC >1.20%, and the mean TOC of the nine samples from Mount Bowers analyzed by Horner and Krissek (1991, their table 5) is 1.31%. In contrast, the mean TOC of Horner and Krissek’s (1991, their table 5) 98 samples from other localities is 0.23%. The high TOC content at Mount Bowers is not due solely to reduced thermal alteration, as indicated by vitrinite reflectance. The mean for five samples from Mount Bowers is 2.78 Ro versus a mean of 2.72 Ro for 31 samples from other localities (Horner and Krissek, 1991, their table 5); the sample with the highest TOC (2.34%) has an Ro of 2.86. Thermal alteration does not affect all organic matter equally. Approximately 10% of TOC may be lost during maturation if it is type III, but up to 80% may be lost if it is type I. If organic matter at Mount Bowers was transported land plant material (type III) rather than algal material (type I), less would have been destroyed by heating, resulting in anomalously high TOC. Higher terrestrial organic input into the Mount Bowers area than elsewhere in the MLIS is consistent with its inferred position near the shoreline; higher plant material is even carried into deep basins of African rift lakes (Talbot, 1988). Macroscopic plant debris should be present if there were a large influx of plant material, but none has been found in the Mackellar Formation, including at Mount Bowers or in the Moore Mountains, which are also interpreted as near the paleoshoreline (Fig. 2). Organic geochemical parameters of samples from the Mackellar Formation do not provide unequivocal evidence about the productivity of the MLIS. High productivity would be demonstrated by high TOC, by kerogen derived from algal sources, and by evidence of anoxia, particularly in the profundal zone. None of these indicators is present, but their absence does not conclusively demonstrate that productivity in the MLIS was low. Organic matter produced in the water column can be degraded by aerobic bacteria. Comparison of algal species in the water column and sediments in modern lakes has verified that some phytoplankton are not preserved in sediments (Livingstone, 1984); in Lake Michigan, only ~2% of the organic matter that is fixed in the epilimnion is preserved in the sediment (Eadie et al., 1992). Bacterial degradation of organic matter would be promoted by oxygen-rich water; the widespread distribution of trace fossils in the Mackellar Formation suggests that the bottom waters of the MLIS were oxygen rich. Amorphous kerogen could be derived from highly degraded algal organic matter, but it is not abundant and is restricted to samples with very low TOC. Although there is no direct evidence regarding nutrient levels in the MLIS, the geologic setting (i.e., its position overlying Permian glacial deposits and Devonian quartz-rich sandstones), and

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absence of palynomorphs and plant macrofossils, suggest a low influx of nutrients. Productivity in the lake might have been much higher than the low TOC would suggest if the organic matter had been altered or destroyed by bacterial degradation in well-oxygenated water or sediments and/or by heating associated with intrusion of Jurassic sills. Available data neither support nor rule out this scenario. The simplest interpretation of the organic geochemical data is that (1) the MLIS received a limited influx of terrestrial organic matter (more in the Mount Bowers area) from streams traversing a recently deglaciated terrain (evidence is low TOC, vitrinite, absence of plant fossils); (2) that the in situ productivity was relatively low, perhaps reflecting low nutrient input; and (3) that the organic matter was affected to varying degrees by bacterial degradation and thermal alteration (evidence is high vitrinite reflectance) caused by Jurassic intrusions.

TABLE 4. CHARACTERISTICS OF MACKELLAR LAKE OR INLAND SEA BASED ON LITHOLOGIC, ICHNOLOGIC, SEDIMENTOLOGIC, AND ORGANIC GEOCHEMICAL DATA FROM THE MACKELLAR FORMATION Setting Cratonic, postglacial, ~80°S Depositional processes

Deposition dominated by turbidity currents. Evidence for waves, storms restricted to rare symmetrical ripples. No hummocky crossstratification

Salinity

Fresh water

Total organic carbon in sediments

Low (mostly <0.5%)

Extent

Outcrop belt ~500 km × 100 km in area studied; correlative units extend an additional 1000 km

Water depth

>20 m–30 m; constrained by thickness of upward-coarsening sequences

LIMNOLOGY OF THE MACKELLAR LAKE OR INLAND SEA (MLIS) Limnologists study modern lakes as systems with complexly interacting biological, physical, and chemical components constrained and controlled by geologic setting and climate. Ancient lake systems increasingly are interpreted as reflecting the interplay between tectonics, climate, and the array of diverse interacting biotic and abiotic factors (e.g., Cohen, 2003; Bohacs et al., 2003; Gierlowski-Kordesch and Kelts, 2000). Important characteristics of lakes required for understanding their functioning as ecosystems include the (1) tectonic, geomorphic, and climatic setting; (2) size, depth, and morphology; (3) mixing regime (e.g., whether or not it is stratified, and, if so, the frequency and extent of mixing); (4) productivity; and (5) oxygen availability. The study of ancient lake deposits, such as the Mackellar Formation, yields information about depositional processes, temporal and spatial distribution of the lake facies and of its fossilized flora and fauna, organic matter preserved in the sediments, and oxygen conditions during deposition. Known characteristics of the MLIS based on lithologic, sedimentologic, ichnologic, and organic geochemical data are summarized in Table 4. The challenge is to use those data to constrain limnologic parameters in order to reconstruct the setting and functioning of the ancient lake ecosystem. Ecologically significant characteristics discussed below include the number of ice-free days; type of stratification and mixing; size, shape, and volume of the lake; and productivity. Ice-Free Days (IFD) The number of days during the year that a lake is not covered by ice controls the temperature regime, light penetration, productivity, and stratification. At the continental scale, the number of IFD on low altitude lakes is controlled primarily by mean air temperature and secondarily by morphometry (fetch and mean depth). A regression equation that predicts the number of IFD based on mean air temperature and mean depth (a proxy for lake

depth) data from 59 North American lakes located from 41° N to 75° N was used to constrain the IFD of the MLIS (Shuter et al., 1983; Kalff, 2002). Climate modeling for the Permian indicates mean summer temperatures at >60° S for both the Sakmarian and Wordian (Early and Middle Permian, respectively) of about –2 °C, and a mean winter temperature of –30 °C for a mean annual temperature of –14 °C (Gibbs et al., 2002). Climate simulation for the southern polar region in the latest Permian (after significant warming had occurred) yields a mean summer temperature of 12 °C and a mean winter temperature of –20 °C for a mean annual temperature of –8 °C (Kiehl and Shields, 2005). The depth of the MLIS is known only to have exceeded the thickness of upward-coarsening sequences (20 m to 30 m). To ensure that MLIS conditions were bracketed, numbers of IFD were calculated for mean annual temperatures of –20 °C to 0 °C and mean lake depths of 20 m to 500 m (Fig. 9; only 12 modern lakes have maximum depths >500 m; Kalff, 2002, table 4.5). The major points gleaned from this estimate are that the MLIS was almost certainly ice covered a significant portion of the year, but also it was likely to have been ice free for at least two months each year, long enough for photosynthesis to occur and the surface water to warm. The most likely range of mean annual temperature based on the Permian climate models is –14 °C to –8 °C. This corresponds to a minimum number of IFD of 80 and a maximum of ~150, depending on lake depth; the deeper the lake, the more IFD (Fig. 9). Stratification and Mixing Stratification and mixing in lakes exert strong control of transport of materials (e.g., nutrients, oxygen) vertically within the lake and thus are important in determining the structure of the biotic communities and lake productivity (Kalff, 2002, p. 154). Light penetrating the upper layers of the lake during the summer

Mackellar Formation, Transantarctic Mountains

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250 20 m 50 m 100 m 200

200 m

Ice-free Days/Year

500 m

Figure 9. Number of ice-free days (IFD) for the MLIS, calculated using different mean annual temperatures and various lake depths that bracket the depth of the MLIS and temperature for Permian high southern latitude (Gibbs et al., 2002; Kiehl and Shields, 2005). Most likely the temperature was –8 °C to –14 °C; the most likely depth was 100 m to 300 m. Estimates are based on the following equation derived from data from 75 modern Northern Hemisphere lakes (Shuter et al., 1983; Kalff, 2002, p. 157): lnIFD = 0.06*(TEMP) + 0.073*(lnzmean) + 5.005.

150

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both promotes photosynthesis by planktic organisms and warms the water, thereby reducing density; postmortem decay of organisms sinking to the bottom depletes oxygen that is not replenished if the water is thermally stratified, leading to low-oxygen conditions in the water column and sediments. Was the MLIS stratified or mixed? Several lines of evidence based on characteristics of modern lakes suggest that it was seasonally stratified. The stability of stratification decreases as the temperature differential between the water at the surface and at depth decreases, a difference that tends to decrease with increasing latitude (Geller, 1992). A simple measure of the likelihood of stratification of temperature in a lake is the ratio of the thickness of the epilimnion, the upper zone of the water column where photosynthesis occurs, to the maximum depth of the lake. If this ratio is <0.5 the lake undergoes stable seasonal (summer) stratification. If it is >2.0 there is no stratification; if it is between 0.5 and 2.0 there may be stratification overturned by strong winds or intermittent stratification (Kalff, 2002, p. 165). The thickness of the epilimnion corresponds to the thickness of the mixed layer and depth to the top of the thermocline. Mixing depth and depth to the thermocline were calculated for the Permian MLIS, using models presented by Kalff (2002, table 11.2) based on data from lakes in Poland and Canada (mixing depth) and lakes worldwide (depth to thermocline); Kalff presented no models based on modern polar lakes. Mixing depth is dependent on the maximum effective length and maximum effective width of the lake, neither of which is known for the MLIS; the depth to the top of the thermocline is based on lake area, which is also

unknown. What is known for the MLIS is the area of outcrop; the region considered here is restricted to the Nimrod, Beardmore, and Shackleton Glacier areas and does not include areas of outcrop of correlative units (Fig. 2). If rocks exposed at Mount Bowers (Fig. 2) were deposited in lake-marginal environments, the dimensions of the outcrop of the Mackellar Formation are ~500 km by 200 km. Used as the maximum effective length and maximum effective width, these numbers correspond to a mixing depth of 48 m (Table 5); even if the effective maximum length and width were much larger than that indicated by the outcrop area, the mixing depth still would be <65 m (Table 5). However, because the age of the Mackellar Formation at different outcrops is poorly constrained, it is possible that the Mackellar facies was deposited in separate small lakes extant at different times in different places along the outcrop belt. The paucity of shoreline deposits (e.g., wave-influenced versus turbidity-current deposits) points to one (or a few) large lakes, but the possibility of a series of small lakes cannot be ruled out. The depth to the thermocline based on areas of an MLIS with the dimensions given in Table 5 is shown in Figure 10. For a lake 500 km by 100 km (100,000 km2) the depth to the thermocline is ~20 m. The major point underscored by Table 5 and Figure 10 is that depth to thermocline and mixing depth are remarkably unchanged for large variations in mean effective length and width and in area. Mixing depth is reduced by half (from ~60 m to ~30 m when mean length is reduced by 90% and mean width by 85%) (Table 5); similarly, depth to thermocline varies little until lake size becomes very small (<100 km2; Fig. 10).

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TABLE 5. MIXING DEPTH (zmix) OF LAKE SURFACE WATER AS A FUNCTION OF MEAN EFFECTIVE LENGTH (MEL) AND MEAN EFFECTIVE WIDTH (MEW) Length (MEL) Width (MEW) Mixing depth (zmix) (km) (km) (m) 1000 200 63 750 200 58 750 150 56 500 200 51 500 150 49 500 100 48 400 100 44 300 100 40 200 100 36 200 50 33 100 50 27 50 50 23 50 20 20 10 2 10 5 1 7 Note: Based on following equation derived from characteristics of 88 extant lakes in Poland and Canada; results show mixing depth remains between 40 m and 60 m for a wide range of lake dimensions: 0.41 zmix = 4.6*([MEL + MEW]/2) .

Using depth to thermocline or mixing depth as a proxy for thickness of the epilimnion (ze) and maximum depths (zm) ranging from 20 m to 500 m, the ze/zm for MLIS consistently is <0.5 for all realistic values. This strongly suggests stable seasonal stratification (Kalff, 2002, p. 165). The MLIS is reconstructed as most likely a dimictic lake that was covered with ice part of the year and thermally stratified during the summer, with two periods of mixing occurring after ice melt and after fall cooling of surface water following warming during the summer. The only uncertainties that arise regarding this interpretation are the applicability of models that are not based on modern polar lakes and whether or not in the MLIS there was a sufficiently large temperature differential between surface and deep water for stratification to occur. Thus, the possibility that the MLIS was a cold monomictic lake cannot be completely excluded, although it is less likely because of the large number of IFD days and the robustness of the model results that point to a dimictic lake. Productivity TOC in lake sediments commonly is interpreted to reflect the amount of biomass that sinks to the bottom, and the lake’s productivity. In oligotrophic (low productivity) systems, TOC in

25

Thermocline Depth, m

20

15

10

5

0 0

20000

40000

60000

80000

100000

120000

140000

160000

180000

200000

Lake area (sq km)

Figure 10. Depth to thermocline (zt) in modern lakes modeled as a function of lake area (A): zt = ln([A/0.5]/0.043)2.35. Note that depth to themocline varies relatively little until lake size becomes very small (e.g., <100 km2; Kalff, 2002).

Mackellar Formation, Transantarctic Mountains sediments provides a reliable qualitative record of productivity under a variety of mixing regimes (Cohen, 2003, p. 253). However, in areas of high productivity, patterns of mixing, bacterial degradation of organic matter in the water column, and dilution of organic matter by clastic sediments are factors controlling the amount of TOC preserved in the sediment. The absence of macroscopic plant fossils in the Mackellar Formation implies that nutrient input from land plants was low (probably too low to support high productivity), and the low level of bioturbation suggests that biologic mixing in the sediment did not enhance bacterial degradation enough to significantly reduce TOC in the sediment. The low TOC, the postglacial setting, the widespread distribution of trace fossils (if low abundance), and the absence of well-preserved allochthonous or autochthonous organic matter are all consistent with the MLIS as an oligotrophic, low productivity lake. CONCLUSIONS: OVERVIEW OF THE MLIS AND ITS FUNCTIONING The MLIS was situated on a recently deglaciated, cratonic, presumably low-relief terrain at southern polar paleolatitude with a cold but not frigid climate. Braided streams traversed an outwash plain before entering the MLIS, loaded with suspended sediment, and formed turbidity currents that transported fine sand, silt, and mud away from the lake margin. The MLIS was covered with ice from a minimum of two months per year to a maximum of six months per year. It was stratified in the summer but was mixed twice yearly; mixing was enhanced by turbidity currents. Bottom waters remained oxic, as indicated by widespread if not pervasive bioturbation. The distribution of bioturbation, a proxy for infauna, indicates that animal activity was least present in areas of rapid sedimentation near the lake margin or adjacent to crevasse splays at turbidite-system channels and was greatest in shallow littoral zones away from loci of deposition and in the distal portions of the lake. Some organic matter was delivered to the lake from the watershed, but the absence of plant macrofossils indicates that vegetation probably was not lush in the recently deglaciated area. Low TOC in the Mackelllar Formation probably accurately reflects a limited contribution of allochthonous organic matter and low productivity within the lake, although productivity in the surface waters might have been higher than suggested by the very low TOC values. The size and depth of the MLIS are poorly constrained. If the size is approximated by the extent of the Mackellar Formation outcrop, the MLIS covered ~50,000 km2 in the present-day Nimrod, Shackleton, and Beardmore Glacier areas. Alternatively, as discussed above, the area may have been covered by a series of small time-transgressive lakes. A major factor favoring a large lake interpretation is the relative lack of shallow-water deposits that would constitute the dominant facies if there had been numerous small lakes. The depth of the MLIS was greater than the ~20 m to 30 m thickness of upward-coarsening sequences. The maximum depth

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is constrained only by comparison with modern glacial lakes, most of which are <500 m deep. Pleistocene glacial Lake Agassiz, the largest proglacial lake known with a surface area of 350,009 km2, had a maximum depth of 200 m (Kalff, 2002). The MLIS resembles glacial Lake Agassiz in its cratonic setting and in the fact that it is underlain by glacial deposits. Recent evaluation of the thickness and extent of the late Paleozic glaciation in Antarctica indicates coverage by a single advance of a smaller and thinner ice sheet than that previously envisioned (Isbell et al., 2003). This ice sheet was unlikely to have scoured to a depth of 500 m; a maximum lake depth of 200 m is more plausible. The relative lack of MLIS deposits indicative of storm or wave activity is perplexing. The fetch of a large lake would be sufficient to produce long-length waves that would rework bottom sediments to a significant depth, but hummocky cross-stratification is not seen, and symmetrical ripples occur at only one locality (Mount Weeks, Fig. 2). The slope of the bottom may have been sufficiently steep that only a small area was above storm wave base, but this is unlikely, given the cratonic setting of the MLIS. Alternatively, lake margin deposits may have been entrained in turbidity currents. It is also possible that major storms occurred when the MLIS was covered by ice; this would have been most likely if the number of IFD was near the upper end of the range suggested by models of the Permian climate (Fig. 9). Although the morphology of the MLIS remains poorly known, integration of sedimentologic, ichnologic, and organic geochemical data with information derived from studies of modern lakes yields a remarkably clear, if incomplete, picture of the functioning of this Permian postglacial lake. Because the record of the MLIS is better preserved by the extensively exposed Mackellar Formation than are other similar, co-eval lakes elsewhere in Gondwana, the reconstruction of the MLIS provides a model of an ancient high latitude lake system that may be applied fruitfully to other less well-exposed Paleozoic postglacial lake deposits. ACKNOWLEDGMENTS We were supported by U.S. National Science Foundation grant ANT 0126146 and OPP 0440954 to Miller and NSF grant ANT 0126086 to Isbell. We thank Elizabeth Gierlowski-Kordesch and Kevin Bohacs for helpful comments on an earlier version of this manuscript. REFERENCES CITED Aitchison, J.C., Bradshaw, M.A., and Newman, J., 1988, Lithofacies and origin of the Buckeye Formation: Late Paleozoic glacial and glaciomarine sediments, Ohio Range, Transantarctic Mountains, Antarctica: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 64, p. 93–104, doi: 10.1016/ 0031-0182(88)90145-9. Archanjo, C.J., Silva, M.G., Castro, J.C., Launeau, P., Trindale, R.I.F., and Macedo, J.W.P., 2006, AMS and grain shape fabric of the late Paleozoic diamictites of the southeastern Parana Basin, Brazil: Journal of the Geological Society [London], v. 163, p. 95–106, doi: 10.1144/0016-764904 -155. Babcock, L.E., Isbell, J.L., Miller, M.F., and Hasiotis, S.T., 2002, New late Paleozoic conchostracan (Crustacea, Branchiopoda) from the Shackleton

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