The Sedimentary Basins of the United States and Canada, Volume 5 (Sedimentary Basins of the World)

VOLUME FIVE SEDIMENTARY BASINS OF THE WORLD THE SEDIMENTARY BASINS OF THE UNITED STATES AND CANADA Editor ANDREW D...

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VOLUME FIVE

SEDIMENTARY BASINS

OF THE

WORLD

THE SEDIMENTARY BASINS OF THE UNITED STATES AND CANADA

Editor

ANDREW D. MIALL Department of Geology, University of Toronto, Toronto, Ontario, Canada

Amsterdam Boston Heidelberg London New York Oxford Paris San Diego San Francisco Singapore Sydney Tokyo

Elsevier Radarweg 29, PO Box 211, 1000 AE Amsterdam, The Netherlands Linacre House, Jordan Hill, Oxford OX2 8DP, UK First edition 2008 Copyright r 2008 Elsevier B.V. All rights reserved No part of this publication may be reproduced, stored in a retrieval system or transmitted in any form or by any means electronic, mechanical, photocopying, recording or otherwise without the prior written permission of the publisher Permissions may be sought directly from Elsevier’s Science & Technology Rights Department in Oxford, UK: phone (+44) (0) 1865 843830; fax (+44) (0) 1865 853333; email: [email protected]. Alternatively you can submit your request online by visiting the Elsevier web site at http://www.elsevier.com/locate/permissions, and selecting Obtaining permission to use Elsevier material Notice No responsibility is assumed by the publisher for any injury and/or damage to persons or property as a matter of products liability, negligence or otherwise, or from any use or operation of any methods, products, instructions or ideas contained in the material herein. Because of rapid advances in the medical sciences, in particular, independent verification of diagnoses and drug dosages should be made British Library Cataloguing in Publication Data A catalogue record for this book is available from the British Library Library of Congress Cataloging-in-Publication Data A catalog record for this book is available from the Library of Congress ISBN: 978-0-444-50425-8 ISSN: 1874-5997 For information on all Elsevier publications visit our website at books.elsevier.com Printed and bound in Hungary 08 09 10 11 12

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CONTRIBUTORS Hugh R. Balkwill Petro-Canada, Calgary, Alberta, Canada Email: [email protected] Benoit Beauchamp Geology and Geophysics, Arctic Institute of North America, University of Calgary, Calgary, Alberta, Canada Email: [email protected] Ronald C. Blakey Geology Department, Northern Arizona University, Flagstaff, Arizona, USA Email: [email protected] Peter M. Burgess Department of Earth Sciences, Royal Holloway University of London, Egham, Surrey, UK Email: [email protected] Octavian Catuneanu Department of Earth and Atmospheric Sciences, University of Alberta, Edmonton, Alberta, Canada Email: [email protected] N. Culshaw Earth Sciences Department, Dalhousie University, Halifax, Nova Scotia, Canada Email: [email protected] J.R. Dietrich Geological Survey of Canada, Calgary, Alberta, Canada Email: [email protected] Jim Dixon Geological Survey of Canada, Calgary, Alberta, Canada Email: [email protected]

Ashton Embry Institute of Sedimentary and Petroleum Geology, Geological Survey of Canada, Calgary, Alberta, Canada Email: [email protected] Frank R. Ettensohn Department of Geological Sciences, University of Kentucky, Lexington, KY, USA Email: [email protected] William E. Galloway Department of Geological Sciences, The University of Texas, Austin, TX, USA Email: [email protected] Martin R. Gibling Earth Sciences Department, Dalhousie University, Halifax, Nova Scotia, Canada Email: [email protected] Raymond V. Ingersoll Department of Earth and Space Sciences, University of California, Los Angeles, CA, USA Email: [email protected] L.S. Lane Geological Survey of Canada, Calgary, Alberta, Canada Email: [email protected] Denis Lavoie Commission Ge´ologique du Canada, CGC-Q/Geological Survey of Canada, GSC-Q Ressources Naturelles Canada/Natural Resources Canada, Que´bec, QC, Canada Email: [email protected]

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Timothy S. Lawton Department of Geological Sciences, New Mexico State University, Las Cruces, New Mexico, USA Email: [email protected] Jock McCracken New Ventures North America, Petro-Canada, Calgary, Alberta, Canada Email: [email protected] D.H. McNeil Geological Survey of Canada, Calgary, Alberta, Canada Email: [email protected] Andrew D. Miall Department of Geology, University of Toronto, Toronto, Ontario, Canada Email: [email protected] V. Pascucci Istituto di Scienze Geologico-Mineralogiche, Universita` di Sassari,

Contributors

Sassari, Italy Email: [email protected] Ryan Post Water Resources Unit, Ministry of the Environment, Hamilton Email: [email protected] Brian D. Ricketts Bbrian D. Ricketts Geological Consulting Ltd., Te Awamutu, New Zealand Email: [email protected] M.C. Rygel Department of Geology, State University of New York College at Potsdam, Potsdam, New York, NY, USA Email: [email protected] Boyan K. Vakarelov Australian School of Petroleum, University of Adelaide, Adelaide, Australia Email: [email protected]

INTRODUCTION

AND

ACKNOWLEDGMENTS

Initial plans for this book were being made in the early 1980s. The series editor, Ken Hsu, approached me to serve as Editor of this volume in 1983, but this would not have been good timing. Commencing in the mid1980s, the Decade of North American Geology (DNAG) project was underway. This endeavour, conceived as a centennial project by the Geological Society of America, and carried out in conjunction with the Geological Survey of Canada (which took responsibility for the Canadian portion of the project), was to result in several shelves full of ‘‘white-covered volumes’’ that represent a quantum leap in our collective knowledge about the North American continent. These books are rich in information, but they were written mainly for the specialist and, with a few exceptions, do not contain accessible summaries or regional syntheses. At the same time, and quite independently, application of the technique of deep crustal seismic reflection profiling was reaping immense benefits in elucidating the structure of the crust down to, and beyond, the level of the Mohorovicˇic´ Discontinuity. Systematic application of the technique was pioneered by Oliver (1980, 1982), leader of the Consortium for Continental Reflection Profiling, which began its work in the late 1970s. In Canada, the comparable Lithoprobe project started its work in 1984, and wrapped up in 2005. DNAG volumes were published throughout the 1990s. My sense was that a synthesis volume about the sedimentary basins of the United States and Canada would be premature until these great projects had been completed, and their lessons absorbed. Accordingly, the start-up of this book was delayed. The first letters of invitation to potential contributing authors went out in 1998. Good writers are busy people, and it took several years, and in some instances, several refusals, before the contents and authorships of the various chapters developed to the point that the eventual production of the book seemed to be possible. Even then, readers will note several major topics that are not covered in the book. In particular, the book does not deal with Alaska or Mexico. Each chapter has undergone several iterations. Each senior contributing author was asked to serve as critical reader for one of the other chapters, the plan being that reading a chapter dealing with a similar topic or adjacent area to the one for which the reader is responsible, might suggest useful comparisons and help to develop crossreferencing between the chapters. Each chapter was also read by myself, and the text of the introductory and concluding ‘‘Postscript’’ chapters were developed in part as analysis and summaries of the contents of the book. In preparing this book we have all been inspired by the work of some of the great synthesizers in our science; individuals who have helped to develop, expand and explain the significance of the plate-tectonics revolution to the study of sedimentary basins, and to the disciplines of stratigraphy and sedimentology. I am thinking particularly of John Dewey, Bert Bally and Bill Dickinson, whose work has been enormously influential over the last 30 years. Larry Sloss and Peter Vail, the founders of modern methods of sequence stratigraphy, also deserve to be recognized in this context. Making use of quantitative models of subsidence and modern seismic-reflection data, and building on the discoveries of COCORP and Lithoprobe, basin analysis is now armed with a superb array of analytical and synthetic tools with which to perfect our understanding of Earth’s continental crust. My own personal thanks go out to such colleagues as Ray Ingersoll, Ash Embry and Nick Eyles, who have consistently helped me to see the big picture and to write more clearly. Ron Blakey’s grasp of evolving paleogeographies, Peter DeCelles’ masterly analyses of foreland basins and their structural evolution, Phil Allen’s rigorous and formal quantitative treatment of basin evolution, have all served to sharpen my own thinking. My experience in North America began in 1965 with a graduate course on the geology of the continent given by David Baird, the then Chair of the Geology Department at University of Ottawa. Amongst his first lessons was the correct pronunciation of ‘‘Newfoundland,’’ but it was only much later that I learned that David had been the Chief Geologist of the Province of Newfoundland, and was instrumental in proposing Gros Morne as a national park on the basis of the fascinating and superbly exposed geology there. I was not to visit this park until Nick Eyles and I were researching our book Canada Rocks: The Geologic Journey along with my wife, Charlene, in 2003. My real introduction to basin analysis came with a seven-year term as Research Scientist with the Geological Survey of Canada in Calgary (1972–1979), where I was privileged to work alongside and learn from such individuals as Hans Trettin, Ray Thorsteinsson, Jim Aitken, Chris Yorath, Don Norris and Don Stott. David Strangway, Geoff Norris and Frank Beales eased my transition to university life in Toronto in 1979.

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Introduction and Acknowledgments

My wife, Charlene, my co-investigator in several projects concerning the social construction of science, and my field companion since a memorable summer in the Arctic in 1976, has been a continuing source of support and good advice. Our children, Chris and Sarah, have learnt to understand Dad’s sometimes odd preoccupation with the rocks and not the scenery on our various world travels. I am grateful for all their love and support. Andrew D. Miall Toronto January 2008

CHAPTER 1

The Phanerozoic Tectonic and Sedimentary Evolution of North America Andrew D. Miall and Ronald C. Blakey

Contents 1. Introduction 2. The Major Phases of Tectonic Development 3. Phase One: The Construction of Pangea 3.1. Plate-tectonic evolution 3.2. Sedimentary evolution of the interior and western continental margin 3.3. Sedimentary evolution of the eastern continental margin 3.4. Sedimentary evolution of the southern margin 3.5. Sedimentary evolution of the Arctic margin 4. Phase Two: Development of the Southern Mid-Continent and Ancestral Rockies 4.1. Plate-tectonic evolution 4.2. Sedimentary evolution of the Mid-Continent and Ancestral Rockies 5. Phase Three: Breakup of Pangea and Formation of the Cordilleran Orogen 5.1. Plate-tectonic evolution 5.2. Sedimentary evolution of the western margin 5.3. Sedimentary evolution of the western interior 5.4. Sedimentary evolution of the Arctic margin 5.5. Sedimentary evolution of the Atlantic and Gulf margins 6. Late Cenozoic Modifications Acknowledgments References

1 2 6 6 13 15 16 16 16 16 17 17 17 23 24 25 25 26 27 27

Abstract The Phanerozoic history of North America can be divided into three broad phases: During the first phase, which lasted from the Late Precambrian to the Triassic, Pangea was under construction. The western continental margin was either a divergent (‘‘passive’’) margin, facing the paleo-Pacific Ocean (Panthalassa) or a backarc basin bordering that ocean, while the eastern margin, beginning in the Middle Ordovician, underwent convergent and collisional tectonism, with the generation of the Appalachian orogen. Phase two, which extended through the Pennsylvanian and Permian, and overlapped in time with the first phase, saw the southwestern margin of the continent affected by oblique-slip displacement between North America and Gondwana (including portions of what are now Mexico and the southwestern United States), with the development of an orogenic highland called the Ancestral Rockies. Phase three, commencing in the Late Triassic or Early Jurassic, corresponds to the Pangea breakup phase, during which North America drifted westwards (relative to a hot-spot reference frame). The eastern continental margin became the modern extensional Atlantic margin, while the western margin underwent accretionary tectonism leading to the assembly of the Cordilleran orogen.

1. Introduction The purpose of this chapter is to offer a succinct review of the Phanerozoic history of the North American continent, in order to provide a framework within which to evaluate the details of the individual chapters that Sedimentary Basins of the World, Volume 5 ISSN 1874-5997, DOI 10.1016/S1874-5997(08)00001-4

r 2008 Elsevier B.V. All rights reserved.

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Andrew D. Miall and Ronald C. Blakey

Figure 1

Basins areas covered by the chapters in this book.

comprise the remainder of the book. An attempt has been made to develop a coherent narrative of the complex kinematic evolution of the continent, in which subsidence and sedimentation in one province can be understood with respect to deformation and uplift elsewhere, as the plate of which North America is a part underwent drift and rotation, and was subjected to plate-margin and intraplate tectonism. This introduction draws on and refers to the constituent chapters in this book (Figure 1), and builds on the excellent summaries of the tectonic setting and basin history of North America prepared for the Decade of North America project, starting with Volume A (Bally et al., 1989; Bally, 1989), plus many other sources, as noted throughout the text.

2. The Major Phases of Tectonic Development The ancient core of North America is the Canadian Shield, which crops out across the northern half of the continent, underlies the cratonic sedimentary cover of the continent’s vast interior plains, and extends in attenuated and metamorphosed form deep beneath the orogens on the continental margins. The Shield was built in three broad stages, which are thought to represent cycles of supercontinent assembly (Figure 2). Around the margins of North America, on the east, south, west, and north, this basement, with its cratonic cover, is stretched, thinned, and buried beneath the complexly deformed rocks of the Appalachian, Ouachitan, Cordilleran, and Innuitian orogens, respectively. These belts are in part accretionary; that is, contained within the orogen are small to large fragments, ranging from tectonic slivers to microcontinents, derived from other, non-North American sources, or from pericratonic sources of probable North American affinity (e.g., Kootenay Terrane in Cordillera).

Phanerozoic Tectonic and Sedimentary Evolution of North America

3

Figure 2 The assembly of the North America continent, in ¢ve broad stages: (1) The original North American continent, Arctica, which started to form about 2.5 Ga; (2) Area added during the formation of Nena, about 1.9 Ga; (3) Grenville orogen, added to complete the formation of Rodinia, between 1.3 and 1.0 Ga; (4) Appalachian orogen, added between 600 and 300 Ma; (5) Cordilleran orogen, added during the breakup of Pangea, commencing about 250 Ma (after Eyles, 2002).

The convergent tectonism which formed the orogens began in the east, during the Ordovician, and continued there until the Permian (area 4 in Figure 2). On the western margin, localized episodes of convergent-margin tectonism have been documented from as far back as the Devonian (Antler orogeny), but the main phase of Cordilleran orogen development began during the Jurassic. From a continental perspective, orogeny and all its consequences (earthquakes, igneous intrusion, volcanism, metamorphism, deformation) has been virtually continuous on the western continental margin since then, with active orogeny extending from Alaska to Mexico (area 5 in Figure 2). It is convenient to commence the narrative in the Late Precambrian, because this is when the ancient and structurally stable central shield achieved the composition and outline shape that it exhibits at the present day. As documented in detail by Hoffman (1988, 1989), the Canadian Shield is the product of a series of collisional orogens that took place between Archean and Late Proterozoic time, the last and one of the most important being the Grenville orogeny, which climaxed at about 1 Ga. This was the final stage in the assembly of the supercontinent now named Rodinia (Dalziel, 1991; Hoffman, 1991). Breakup of Rodinia commenced with widespread rifting at about 800 Ma, culminating in the isolation of North America as a separate continent by about 600 Ma (Chapters 3 and 4). During the period when the plate was undergoing rifting and separation from the rest of Rodinia, and leading up to its involvement in the assembly

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Figure 3

Andrew D. Miall and Ronald C. Blakey

The Laurentian plate at the end of the Precambrian (after Ho¡man, 1989).

of Pangea, the continent is referred to as Laurentia. By the end of the Precambrian, Laurentia was almost entirely rimmed by extensional margins (Figure 3), much as are Africa and Antarctic today following the breakup of Pangea (Bally, 1989), and for much the same reason — the continent was surrounded by sea-floor spreading ridges from which other continents rotated away (Hoffman, 1988, 1989, 1991). The Late Precambrian to Early Paleozoic history of Laurentia is characterized by thick sedimentary successions in rift basins and extensional-margin sedimentary wedges (Figure 3). Several ‘‘failed rifts’’ extend into the continental interior, some of which became substantial basins during the Late Precambrian and Paleozoic (e.g., Southern Oklahoma rift; Chapter 8). The crustal fragmentation that occurred during this period continued to provide the focus for intraplate earthquakes throughout the Phanerozoic. The Reelfoot Rift beneath the Mississippi valley and the St. Lawrence Rift of eastern Canada are particularly well known as the loci of modest to major earthquakes. The Phanerozoic history of North America can be divided into three broad phases: (1) During the first phase, which lasted from the Late Precambrian to the Triassic, Pangea was under construction. The western continental margin was either a divergent (‘‘passive’’) margin, facing the paleoPacific Ocean (Panthalassa) (Chapter 5), or a backarc basin bordering that ocean, while the eastern margin, beginning in the Middle Ordovician, underwent convergent and collisional tectonism, with the generation of the Appalachian orogen (area 4 in Figure 2; Chapters 3 and 4). Baltica and Africa continued to move against eastern North America until the Permian, generating significant extensional and strike-slip displacements on the eastern margin in some places (e.g., parts of Atlantic Canada; Chapter 6), but also significant fold-andthrust belt development, as in the Appalachian Basin and Ouachita foreland basin (Chapter 8). (2) Phase two, which extended through the Pennsylvanian and Permian, and overlapped in time with the first phase, saw the southwestern margin of the continent affected by oblique-slip displacement between North America and Gondwana (including portions of what are now Mexico and the southwestern United States),

Phanerozoic Tectonic and Sedimentary Evolution of North America

5

with the development of an orogenic highland called the Ancestral Rockies (Chapter 7). This phase is defined primarily for the significantly distinct history that the southwest part of the continent experienced because of the unique regional plate kinematics. (3) Phase three, commencing in the Late Triassic or Early Jurassic, corresponds to the Pangea breakup phase, during which North America drifted westwards (relative to a hot-spot reference frame: Engebretson et al., 1985). The eastern continental margin became the modern extensional Atlantic margin (Chapter 14), while the western margin underwent accretionary tectonism leading to the assembly of the Cordilleran orogen (area 5 in Figure 2; Chapters 10 and 11). Events on the northern (Arctic) margin of North America can be correlated to phases 1 and 3, as discussed further (Chapter 13). A major contrast between phases 1 and 3 is that, until the Triassic, North America was undergoing convergence with Africa-Europe in the east, while a long-lived extensional margin faced Panthalassa in the west, and then from Late Triassic time to the present, relative motions reversed, with an extensional margin facing the widening Atlantic Ocean in the east, while a complex, convergent, accretionary orogen developed in the west. This evolution is encapsulated by the series of cross-sections across Canada shown in Figure 4.

Figure 4 Sequential east-west cross-sections through Canada, summarizing the changes in the plate-tectonic regime during the Phanerozoic.

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Andrew D. Miall and Ronald C. Blakey

3. Phase One: The Construction of Pangea 3.1. Plate-tectonic evolution During the Cambrian, Laurentia lay astride the equator, rotated about 901 clockwise, relative to its present-day orientation (Figure 5). The Iapetus Ocean along Laurentia’s eastern margin functioned as a growing ocean for about 100 Myr (this is less than half the duration of the present-day Atlantic Ocean), and is estimated to have been about 5,000 km wide by the end of the Cambrian. The divergent margin was marked by a series of promontories and embayments, formed by transform offsets in the Iapetus spreading center. The largest of these are the Newfoundland promontory with the Quebec embayment immediately to the south, and the Alabama promontory and Ouachita embayment (Thomas, 2006; Figure 6). These promontories had a pronounced effect on tectonism and the development of sedimentary basins during the subsequent series of Appalachian orogenies (Ettensohn, Chapter 4). By Mid-Cambrian time, closure of Iapetus Ocean east of the continent (south, in Cambrian time) had commenced, with the establishment of at least one arc complex offshore (TAC in Figure 5). The collision with and deformation of these arcs against Laurentia initiated the Taconian Orogeny. This was the first in a series of orogenies that rafted fragments of Gondwana against the eastern Laurentian margin, resulting in a complex orogenic collage that has taken some twenty years to fully comprehend (from Williams, 1978, to Van Staal et al., 1998). Traditionally, the Taconian orogeny has been interpreted to have begun in the Middle Ordovician; however, some evidence (see Chapter 3) indicates that it may locally have commenced in the latest Cambrian. During the Ordovician, the Taconian orogeny affected the entire eastern margin of the continent (Figure 7), and formation of the Appalachian foreland basin commenced (Figure 8; Lavoie, Chapter 3; Ettensohn, Chapter 4). At about this time a new ocean, variously called the proto-Tethys Ocean or Rheic Ocean, formed between Baltica and Africa (Figures 9 and 10). Rifted fragments of Africa and South America were rafted by seafloor spreading within this ocean, and eventually collided with and sutured against Laurentia. A fragment interpreted as part of the continental margin of what is now Morocco — The Gander Terrane — collided with Laurentia during the Early Silurian, simultaneously all along the continental margin, from Newfoundland to Maine, causing what has been termed the Salinic Orogeny (Chapters 3 and 4). Eastern Newfoundland, and parts of Nova Scotia, New Brunswick, and Maine (and parts of central England) are now composed of the Avalon terrane. The arrival and docking of this microcontinent during the Silurian was the cause of the Acadian Orogeny, which has been well documented in New Brunswick and Maine (Chapters 3 and 4). This was one of the last events in the long series of arc–continent and continent–continent collisions which ultimately brought Laurentia and Baltica together. European geologists call this collision the Caledonian orogeny (Caledonia is the ancient name for Scotland), based on the evidence for Ordovician to Early Devonian tectonic episodes along what is now the suture between these two continents. A recent geodynamic model proposed for Appalachian, Caledonian, and Early Variscan terrane accretion by Stampfli et al. (2002) greatly simplifies plate-tectonic events and reduces the number of accreted plates. They proposed that two major ribbon-like terranes rifted from northwest Gondwana during the Early Paleozoic and rafted across the Iapetus and Rheic-proto-Tethys oceans. The first of these terranes, Avalonia (generally divided into East and West — EAV and WAV in Figures 9 and 10), docked with Baltica and northeastern North America commencing in the Early Silurian — coeval with the collision of Baltica and Laurentia (Figure 9). This triple collision closed the Iapetus Ocean and generated the Caledonian orogeny, sensu stricto of McKerrow et al. (2000). The Gander terrane of New England and Atlantic Canada may have been part of this accreted terrane (the Salinic orogeny, noted above). West Avalonia scissored into southern New England and the Mid-Atlantic states in the Late Silurian and Early Devonian (early phase of Acadian orogeny; Figure 10). Stampfli et al. (2002) placed Meguma on the trailing margin of the Avalonia plate. The second ribbon terrane, the Hunic terrane (HUN in Figure 11), closed the Rheic Ocean and diachronously accreted to Laurentia and southern Europe commencing in the Late Devonian to the northeast along East Avalonia, and in the Mississippian-Pennsylvanian to the southwest along the Gulf Coast states (Figure 11; Early Variscan, Late Acadian, and Ouachita–Marathon orogenies, respectively; see Chapters 6 and 8). Accreted blocks of the Hunic terrane include the Ligerian, Amorican, and Iberian terranes of Europe, the Meseta terrane (now in North Africa), and the Carolina, Yucatan, Oaxacan, and Chortis terranes of Mexico and Central America. Whereas previous work had suggested numerous terrane-accretion events to explain the complex juxtaposition of plates, Stampfli et al. (2002) postulated transpressional forces as responsible for doubling and tripling the ribbon terranes (especially in the southern Iberian peninsula and southwestern France; see their Figure 4) during complex lateral folding and shearing in the ensuing collision of Gondwana with Laurussia (Figure 12). By the end of the Silurian and the beginning of the Devonian, Iapetus had closed everywhere. A major tectonic episode, termed the Scandian phase of the Caledonian Orogeny, occurred when Baltica collided with

Figure 5 Plate-tectonic setting of North America during the Middle Cambrian. These and subsequent maps have drawn extensively on the data and ideas contained in Cook and Bally (1975), Ziegler (1988), Scotese (1998), Scotese and Golonka (1992), and Stamp£i et al. (2002). AFR, Africa; ANT, Antler terrane; ARC, Artica; ARM, Ancestral Rocky Mountains; BAL, Baltica [N. Europe]; EAV, East Avalonia [England]; FLA, Florida; GRN, Greenland; GUE, Guerro [W. Mex]; HUN, Hunic terranes [central, south Europe, SE US, Mexico, C. America]; MEG, Meguma [E. US, part of Avalonia?]; NAM, North America [ ¼ Laurentia]; NSL, North Slope [Alaska]; PRE, Precordillera; QUE, Quesnell; SAM, South America; SCT, Scotland; STI, Stikine; TAC,Taconia [E. US]; WAV,West Avalonia [E. US];WRG,Wrangellia; YUC,Yucatan. Additional maps and references may be found at http://www4.nau.edu/geology/.

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Figure 6 The eastern North American margin, during the Early Cambrian. At this time eastern North America was functioning as an extensional margin, bordering the Iapetus Ocean (Thomas, 2006).

Greenland (Figure 10). The reverberations of this episode were felt in the Arctic Islands as a series of small uplifts in the central and eastern Arctic. The Meguma terrane (MEG in Figures 9 and 10) rifted from North Africa and arrived off Nova Scotia in the Mid- to Late Devonian. The suturing of northern Laurentia and Baltica formed an orogenic highland at the center of this large combined continent — Laurussia — which has also been called the ‘‘Old Red Continent’’ because it was the source for and the site of the predominantly coarse, red clastics constituting the Old Red Sandstone of western Europe, Svalbard, and Atlantic Canada. Detritus was also shed westward across Greenland into the Canadian Arctic, and probably contributed to clastic-wedge formation as far west as Yukon Territory. Between the Mississippian and the Permian, North America underwent orogenic collision on three of its four margins (Figures 11–13), and by the end of this period North America had become one of the larger components of the new supercontinent Pangea. During the Late Devonian or Mississippian, the east-facing Antler arc approached the western margin, and collided with the Cordilleran miogeocline (Ingersoll, Chapter 11; Figure 11). Distal, passive-margin, deep-water deposits were thrust eastward over coeval carbonate shelf deposits as the Roberts Mountains Allochthon. The arc itself collided with partly rifted fragments off Western North America; the largest blocks became the nucleus for the Quesnell and Stikine terranes, which were accreted to North America in the Mesozoic. During the Carboniferous to Early Triassic, these terranes were approaching North America from the west. Active arcs, within which significant volumes of Panthalassa oceanic crust were being subducted, lay off the western continental margin (Figure 13). These arcs and associated terranes began to accrete to the western margin in the Permian and Triassic (Sonoma orogeny of California–Nevada—Oregon, Chapter 10). Until the Late Triassic, much of the continental margin off western Canada remained in an extensional-margin tectonic and sedimentary (‘‘miogeoclinal’’) regime, with the accumulation of craton-derived shallow- to deep-marine deposits, including turbidites (Chapter 5).

Phanerozoic Tectonic and Sedimentary Evolution of North America

Figure 7

Plate tectonic setting of North America during the Middle Ordovician.

Figure 8

Clastic wedges that developed as a result of the Appalachian orogenies (Thomas, 2006).

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Figure 9

Figure 10

Andrew D. Miall and Ronald C. Blakey

Plate-tectonic setting of North America during the Early Silurian.

Plate-tectonic setting of North America during the Early Devonian.

Phanerozoic Tectonic and Sedimentary Evolution of North America

Figure 11 Plate-tectonic setting of North America during the Late Mississippian.

Figure 12

Plate-tectonic setting of North America during the Late Pennsylvanian.

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Figure 13 Plate-tectonic setting of North America during the Early Permian.

In the east and south, the Gondwana margin (now comprising the west coast of Africa and the north coast of South America) underwent the final collision with North America, causing the Pennsylvanian-Permian Alleghanian orogeny from Alabama to the Atlantic provinces (Chapters 3 and 4) and the Marathon–Ouachita orogeny from Alabama to Texas and into northern Mexico (Figure 12; Chapter 8). The Carboniferous Ouachita basin has been interpreted by Graham et al. (1975) and Ingersoll et al. (1995) as a remnant ocean basin between the approaching Laurentian and Gondwana continents. However, their original model of diachronous collision of Laurentia with Gondwana needs to be modified in light of modern data relating to the importance of terraneaccretion events. The Ouachitan orogeny probably began as a terrane-accretion event in Alabama in the Mid- to Late Mississippian (Figure 11), reaching Mexico in the Early Permian. There would therefore have been a remnant ocean across southern Texas during much of the Pennsylvanian (Chapter 8). There may also have been a remnant ocean behind (south of) these terranes, corresponding to a much reduced proto-Tethys, as suggested in Figure 11. As noted below (under Phase two), this collision was the cause of the uplift of the Ancestral Rocky Mountains in the southwest part of the continent (Chapter 7). The Alleghanian orogeny represents the final collision between Gondwana and Laurentia. It probably commenced in the Early Pennyslvanian, and may have proceeded from south to north, reflecting diachronous contact between irregular continental margins (Ettensohn, Chapter 4). Taconic, Acadian, and Alleghanian orogenies together accounted for 100–400 km of crustal shortening and overthrusting above the basal de´collement beneath the Appalachian mountains of the United States (Ettensohn, Chapter 4). The final stage of Pangea construction, during the Pennsylvanian and Permian, involved strike-slip displacement between Gondwana and Laurussia, with the formation of linear, strike-slip, and extensional basins bordered by fault-bounded uplifts extending through the Canadian portion of the Appalachian orogen. Oblique convergence and diachronous terrane collision were accompanied by crustal delamination, heating of the overlying crust, the generation of large volumes of mafic magma, and also of large volumes of late- to postorogenic felsic magma (Gibling et al., Chapter 6). The cessation of this activity was followed by significant thermal subsidence during the Permian, which accounts for the final configuration of the Maritimes Basin and is probably largely responsible for the development of the present-day Gulf of St. Lawrence. The plate-tectonic evolution of the Arctic margin is less well understood. Paleogeographic considerations suggest the presence of a land barrier and/or sediment source, called Crockerland by Embry and Beauchamp (Chapter 13) north of the present Arctic margin during the Late Paleozoic, and possibly until the Early Jurassic.

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This land may have been a fragment of what became part of Siberia (ARC in Figures 9 and 10), after the Canada Basin opened in the Cretaceous.

3.2. Sedimentary evolution of the interior and western continental margin The Cordilleran margin, from Alaska to California, is characterized by thick prisms of Mid- to Upper Proterozoic sedimentary rocks, including the well-known Windermere, Miette, Belt and Purcell supergroups of Montana-Idaho-Alberta-British Columbia (Hoffman, 1989; Figure 14). Subsidence analysis of the Phanerozoic margin of southern British Columbia indicates that the extensional subsidence which initiated the opening of Panthalassa began at about 600 Ma (Bond and Kominz, 1984; Chapter 5), so these Proterozoic successions are thought to represent several discrete episodes of rifting that preceded the breakup of Rodinia. They are at least as

Figure 14

Proterozoic strata of the Cordilleran margin (Ho¡man, 1989, Figure 48).

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old as 800 Ma, indicating that rifting lasted for some 200 Myr during the breakup phase. This compares with the lengthy pre-drift rifting of the North Atlantic margins. For example, the rifting of East Greenland lasted for 170 Myr before oceanic crust began to be generated between Greenland and Europe (Surlyk et al., 1981). Upper Proterozoic strata exceed 4 km in thickness in parts of British Columbia, and 7 km in Utah (Hoffman, 1989). They may have been derived in part from erosion of the mountains formed by the Grenville collision and transported across Laurentia by a continental drainage system (Rainbird et al., 1997). The strata consist primarily of thick successions of coarse clastics, including coarse, poorly sorted conglomerates, much of it interpreted as glacigenic in origin. Although Upper Proterozoic glacigenic strata have been interpreted in terms of a frozen earth, as part of the Snowball-Earth hypothesis (Hoffman et al., 1998), the presence of thick successions of debrisflow deposits (e.g., McMechan, 2000), while consistent with deposition under temperate glacial, possibly Alpineglacial conditions, indicates the presence of abundant liquid water and the predominance of aqueous environments, which is inconsistent with the frozen-ocean concept that is the central element of the Snowball-Earth hypothesis (see critique and alternative model for Late Precambrian glaciation by Eyles and Januszczak, 2004). Strata of Cambrian to Early Jurassic in age constitute the oldest four of Sloss’s classic Phanerozoic sequences, the Sauk, Tippecanoe, Kaskasia, and Absaroka (Sloss, 1963, 1988; Burgess, Chapter 2; Miall, Chapter 5; Figure 15). The first three, at least, are widely distributed across the interior of the northern part of the continent. It seems likely that during parts of the Ordovician and Silurian, exceptionally high sea levels caused transgressions to cover most, if not all, of the cratonic interior of Laurentia, including most of the area now exposed as the Canadian Shield (Burgess, Chapter 2). Transgression and regression were caused by a combination of eustatic sea-level changes and epeirogenic tilting and warping driven by mantle thermal processes, the latter in part a consequence of collisional orogeny on the continental margins (Burgess, Chapter 2). Successions of largely Early to Middle Paleozoic age constitute the fill of three of the four large intracratonic basins within the continent, the Michigan, Illinois and Hudson Bay basin (The fourth, the Williston Basin, was markedly affected by Mesozoic–Cenozoic sedimentation during the third phase of continental development, when it received sediment from the uplift and erosion of the Cordilleran orogen). Within the Canadian portion of the western Laurentian margin, Cambrian to Triassic sedimentary rocks constitute a structurally relatively conformable succession of continental-margin strata, comprising a classic ‘‘miogeocline’’ (Figure 16; Chapter 5). Areas where section is locally missing (e.g., West Alberta Ridge: Chapter 5) were uplifted as a result of intraplate stresses or possibly as a result of mantle underplating (see Chapter 17). In western Canada and the United States, the strata can readily be subdivided into the Sloss sequences on the basis of regional low-angle unconformities, and they show that strata of Ordovician and Devonian–Mississippian age are the most widespread of the Paleozoic assemblages. Rocks of this age range typically lap onto the Canadian Shield throughout western and Arctic Canada.

Figure 15 The six classic sequences of Sloss (1963).

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Figure 16 Generalized stratigraphic cross-section through the Western Canada Sedimentary Basin (after Price et al., 1972).

Distribution of the Absaroka sequence, of Late Mississippian to Early Jurassic age, is somewhat different from that of the older sequences. Within Canada it is much less areally extensive than the preceding sequences (Sloss, 1988). It is absent from the central and northern cratonic interior, but is well represented in the southwestern part of the continent (Chapters 7 and 8). The sequence rests on a profound unconformity that records extensive preLate Mississippian denudation, and is also capped by an extensive unconformity. These broad characteristics reflect uplift of the continent in response to orogeny along three of the four continental margins (Alleghanian, Ouachita, and Sonoma orogenies; Figures 11–13). During this period Pangea was undergoing final assembly (Chapters 6 and 8). This was a period of low global sea levels, in part because the crust of the supercontinent was thickened and thermally elevated, and possibly also in part because eustatic sea levels were low because of a slowing of global average rates of sea-floor spreading, and a consequent increase in the global average age (and depth) of oceanic crust (e.g., Worsley et al., 1986). Studies of detrital zircons in the clastic units of the US southwest have suggested that much of the detritus there may have been derived by erosion of the Appalachian and Grenville orogens in the east, and transport west by continental river systems (Blakey, Chapter 7; see also Chapter 17).

3.3. Sedimentary evolution of the eastern continental margin Remnants of Iapetus Ocean are now preserved in the Dunnage zone of Newfoundland. These include some pristine fragments of ancient ocean floor, plus some very altered remains that are very difficult to interpret

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(Van Staal et al., 1998). Recent work, discussed briefly in Chapter 17, is revealing a complex of small oceans, arcs, and microcontinents that formed and were rapidly destroyed during the Taconic orogeny. The Iapetan continental margin of North America is spectacularly exposed in Gros Morne Park of Newfoundland, including the thick carbonate debris flows of the Cow Head Breccia (Lavoie, Chapter 3). Remnants of the sedimentary prism are present at many places to the south, along the ancient continental margin at the edge of the Appalachian orogen, notably in the Valley and Ridge province of the Appalachian orogen. There, extending from Pennsylvania to Georgia, a miogeoclinal transition can be traced in the Cambrian and Ordovician succession, from a largely carbonate cratonic and continental-margin succession in the west to a deeper water carbonate-clastic succession in the east (Rast, 1989; Ettensohn, Chapter 4, Figure 14). Foreland-basin clastic wedges associated with the Taconian, Acadian, and Alleghanian orogenies have long been known in the Appalachian orogen (King, 1977; Ettensohn, Chapter 4). Each appears to radiate from a source area of limited lateral extent (Figure 8), and is diachronous along strike. Thomas (2006) and Ettensohn (Chapter 4) interpret this as the result of diachronous collision of the approaching terranes with the irregular Iapetan margin. Taconic collision and crustal loading commenced first against the St. Lawrence promontory, next against the Alabama promontory, followed by progradation into the Tennessee embayment, then against the New York promontory, followed by migration into the Pennsylvania embayment. Strike-slip and extensional basins of Mississippian to Permian age in Maritime Canada and Newfoundland are characterized mainly by continental and shallow-marine clastics, including significant coal deposits, and local thick evaporites (Gibling et al., Chapter 6).

3.4. Sedimentary evolution of the southern margin Along the Ouachita margin, a cratonic carbonate shelf persisted until Mississippian time. Tectonic loading commenced in the east in the Late Mississippian, and extended westward during the Pennsylvanian (Figure 12; Arbenz, 1989; Thomas, 2006; Chapter 8), reflecting the last, diachronous, stages of closure of Gondwana against Laurentia. A Cambrian to Lower Mississippian pre-orogenic, largely clastic succession up to about 3.5 km thick, including turbidites, is present along the Ouachitan margin, with paleocurrent and petrographic data indicating derivation from the craton to the north. This is followed by 12–15 km of largely deep-water turbidite deposits indicating sediment transport along the axis of the remnant ocean basin in a generally west to northwest direction.

3.5. Sedimentary evolution of the Arctic margin At least 10 km of Cambrian to Devonian strata are present in the Franklinian Basin (Trettin, 1989). Although the origins of the basin are obscure, it probably was initiated by crustal stretching and possibly rift faulting during the breakup of Rodinia toward the end of the Precambrian. A central, deep basin extends from northeast Ellesmere Island southwestwards, at least as far as northwest Melville Island, but this part of the basin is largely covered by younger strata. The deep basin is extensively exposed only in Ellesmere Island as a result of uplift during periods of tectonism during the Late Paleozoic and the Mid-Cenozoic. There, a transition from deep-water to shallowshelf environments is well exposed. A similar deep basin and basin-to-shelf transition are preserved to the northeast, in northern Greenland. The shelf is well exposed along a belt extending from northern Ellesmere Island southwestward to Devon and Cornwallis Islands and then westward to Bathurst and Melville Islands. Subsurface data show that the belt continues on southwestward beneath Prince Patrick and Banks Islands. Pearya, a small but complex terrane, collided with the Arctic margin in the Early Silurian. This area recorded the effects of the Caledonian orogeny, with local uplift and synorogenic sedimentation in the earliest Devonian and again in the Mid- to Late Devonian, when a major clastic wedge prograded southwestward from sediment sources within the Caledonian orogen of eastern Greenland. Franklinian Basin sedimentation was brought to an end by uplift at the end of the Devonian.

4. Phase Two: Development of the Southern Mid-Continent and Ancestral Rockies 4.1. Plate-tectonic evolution The final closure of Gondwana against Laurentia during the Late Paleozoic subjected the southwestern part of Laurentia to a transpressive tectonic regime between the Late Mississippian and Early Permian (Figures 11–13). Sedimentation and tectonism of the Ouachita-Marathon belt at this time have been referred to above. The southern portion of the vast, interior cratonic (largely carbonate) platform of the Early to Mid-Paleozoic was

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transformed, beginning in the Late Mississippian, by warping and faulting. Deformation of the craton west of the Mississippi River by transpressive tectonism differentiated this region into a series of basins and uplifts, of which the best known are the Delaware Basin, Central Basin Platform, Midland Basin, plus the Ouachitan foreland basins (Val de Verde, Fort Worth, Anadarko basins: Chapter 8). Tectonism also had a significant effect on Texas and the Four Corners states, where the Ancestral Rocky Mountains developed (Kluth and Coney, 1981; Blakey, Chapter 7; ARM in Figures 12 and 13). Phase-two tectonism is first in evidence during the Late Mississippian, in the Anadarko basin, along the TexasOklahoma border area. This basin underwent particularly rapid subsidence, accompanied by uplift of the adjacent craton, including movement on such trends as the Wichita-Amarillo axis. However, tectonism was most intense during the Middle Pennsylvanian. At this time, the Ancestral Rockies developed across the entire southwestern portion of what is now the United States (from Kansas to Arizona and Idaho to Texas), consisting of broad, block uplifts bounded by narrow fault zones. The final construction of Pangea during the Pennsylvanian and Permian included the Variscan (Hercynian) orogeny between Europe and Africa (Chapter 6), and led to widespread orogenic uplift of eastern North America and western Europe. Burgess (Chapter 2) attributes this to dynamic topographic uplift over a thermal high caused by supercontinent insulation of the mantle. A reorganization of global plate regimes commenced in the Early Carboniferous, as the final construction of Pangea was underway (Ziegler, 1988). Extensional successor basins developed over some of the areas affected by the Caledonian orogeny. The Sverdrup Basin, in Canada’s Arctic, is one of these (Embry and Beauchamp, Chapter 13). Global sea levels were at an all-time low during the Permian-Triassic, and remained low until widespread rifting and fragmentation of Pangea commenced in the Jurassic.

4.2. Sedimentary evolution of the Mid-Continent and Ancestral Rockies Rocks of Pennsylvanian to Permian age (Absaroka sequence) underlie much of the US Mid-Continent region, west of the Mississippi River. Within the craton, from the Dakotas, south to Kansas, and in the intracratonic Illinois basin, they are characterized by the distinctive cyclic repetitions of the classic cyclothems, which were first described from outcrops in Kansas in the 1930s (Chapter 8). Reef carbonates and evaporites were deposited in the classic Permian Basin of west Texas. The margin of the craton at this time lay across central Texas, and the sedimentary succession shows a transition through platform-margin facies into the deep-water deposits of the Ouachita foreland basin to the south (Chapter 8). Uplift of Pangea, centering on the Appalachian–Caledonian megasuture, imposed a tilt to the North American continent, downward to the west, upon which was established a major, west-flowing drainage pattern. Late Paleozoic and Mesozoic strata in the basins associated with the Ancestral Rockies of the southwestern United States are consequently rich in Grenville and Appalachian detritus (provenance and paleogeographic analysis by Dickinson, 1988; Dickinson and Gehrels, 2003). The Ancestral Rockies were characterized by widespread continental sedimentation, including substantial volumes of fluvial and eolian clastics (Blakey, Chapter 7). One of the more prominent of the active tectonic highs was the Uncompahgre Uplift, which extended from northwestern New Mexico, northwestward across Colorado into Utah. This axis is bounded to the southwest by the Paradox basin, and together they exemplify the pattern of basins and uplifts that characterized the southwest through the Mid-Pennsylvanian to Early Permian. Downthrow on the southwest side of the uplift exceeds 3 km, and banked against the fault is a comparable thickness of coarse, nonmarine clastics of the Cutler Formation, derived by uplift and erosion along the Uncompahgre axis. The adjacent Paradox Basin is filled primarily by a carbonate-evaporite succession recording repeated, cyclic changes in salinity in response to high-frequency changes in sea level (Blakey, Chapter 7).

5. Phase Three: Breakup of Pangea and Formation of the Cordilleran Orogen 5.1. Plate-tectonic evolution The western margin of Canada may have largely remained an extensional margin until the Triassic, whereas, to the south, subduction of Panthalassa probably began in the Mid-Devonian, this being the age of arc-related rocks in Nevada and California. Collision of an interpreted east-facing arc in this area gave rise to the Antler orogeny (Oldow et al., 1989; Ingersoll, Chapter 11; Figure 10). Evidence for convergent-margin tectonism elsewhere along the western margin of Pangea is sketchy, and this margin may not have been significantly affected by subduction and arc collision until the Permo-Triassic Sonoma orogeny of Nevada, which initiated the widespread tectonism of the Cordilleran orogen. The Sonoman orogeny developed as the Sonoma terrane collided with the

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extensional margin of North America along an east-facing arc. The Golconda subduction complex developed above (to the east of) this arc (Ingersoll, Chapter 11). By Late Triassic time the west-facing Nicola arc was well established on the west coast, extending roughly north northwest-south southeast through the present location of Kamloops, British Columbia (Oldow et al., 1989; Yorath, 1991; Price and Monger, 2003; Ricketts, Chapter 10) and continuing southward into the United States (Ingersoll, Chapter 11). At this stage, it was an extensional arc, with a backarc basin now preserved in Canada as the Slide Mountain terrane. The Cache Creek Complex, which extends almost continuously through most of British Columbia, is the well-exposed subduction complex of the Nicola arc, containing rocks and fossils as old as Pennsylvanian. It ranges up to Late Triassic in age, and represents one of the major sites where Panthalassa was subducted as North America drifted westward. According to Engebretson et al. (1985), some 13,000 km of Panthalassa oceanic crust have been subducted beneath North America between the Early Jurassic and the present day (a width equal to one-third of the Earth’s circumference). As Price and Monger (2003, p. 21) pointed out, this documented 180-Myr record only represents about half of the probable duration of the period of subduction of Panthalassa, which began at least as far back as the establishment of the Antler arc in the Mid-Devonian. The arc may be even older than this. Oldow et al. (1989, p. 159) noted that the arc rocks rest on a Cambrian–Silurian basement in the Klamath Mountains and Sierra Nevada. The Cache Creek subduction complex, which became inactive following the accretion of Stikinia, represents a site of oceanic subduction in addition to that computed by Engebretson et al. (1985). Most of the Jurassic to Recent subduction would have been at locations farther to the west, as terrane accretion took place and subduction shifted outboard. The Triassic was a time of global plate-tectonic transition (although precursors of this transition had appeared during the Carboniferous, as noted above). An enormous series of rifts developed along what was to become the Gulf and Atlantic borderlands, extending from the Gulf Coast and Florida to Newfoundland, to northern Greenland, and deep into west Europe, as far east as Poland (Ziegler, 1988; Ettensohn, Chapter 4; Miall et al., Chapter 14; Galloway, Chapter 15; Figure 17). Widespread continental sedimentation continued in the southwest, over the eroded roots of the Ancestral Rockies (Blakey, Chapter 7). On the western margin of North America, thin-bedded turbidites exposed near Banff, Alberta (Triassic Spray River Formation) constitute the last major craton-derived sedimentary units before Cordilleran uplift generated western sediment sources and

Figure 17

Plate tectonic setting of North America during the Triassic.

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reversed sediment transport directions. They were deposited on the miogeoclinal eastern flank of the Slide Mountain backarc basin, but have been tectonically transported far to the east by Cordilleran thrust faulting. Although rifting of Pangea commenced in the Triassic (along fractures that, in some cases, began to form in the Permian; Chapter 6), the breakup of the supercontinent and the appearance of the first Atlantic oceanic crust did not take place until the Middle Jurassic. The oldest part of the Atlantic Ocean is that located off the east coast of the United States (Ziegler, 1988; Sheridan, 1989; Miall et al., Chapter 14; Figure 18). The Gulf of Mexico basin is slightly younger, oceanic crust there spanning the Middle or Early Late Jurassic to Early Cretaceous (Worrall and Snelson, 1989; Galloway, Chapter 15). Rifting of the north Atlantic region commenced in the Early Cretaceous, extending northward as far as Svalbard. Separation of Greenland from Canada also began, with rifting initiating the Labrador Sea-Baffin Bay seaway. However, continental breakup and sea-floor spreading did not begin until the Mid-Cretaceous, when the Iberian Peninsula separated from the Grand Banks (Figure 19; Chapter 14). In the Late Cretaceous, Atlantic opening extended into the area between Greenland and Britain. Stretching of the Grand Banks switched direction from east-west to northeast-southwest. The succession of rifting episodes with contrasting stretching directions was the key to the structural evolution of what became the Jeanne d’Arc Basin and the Hibernia oil field (Tankard and Welsink, 1987; Miall et al., Chapter 14). In the latest Cretaceous, sea-floor spreading commenced off the margins of Labrador, separating Canada from Greenland for the first time. A triple-point junction developed off the southern tip of Greenland, and for about 40 Myr, from the Late Cretaceous until the Oligocene, Greenland functioned as a separate plate (Srivastava et al., 1981). Greenland rotated away from the Labrador–Baffin Bay margin around a pole that caused it to contract against Ellesmere Island, and this had important consequences for the development of the mountains of the northeastern Arctic Islands. The Late Cretaceous–Mid-Tertiary Eurekan orogeny was an episode of transpressive deformation of the northeastern Canadian Arctic Islands. Between the Early Jurassic and the present day, North America drifted westward by some 701 of longitude, relative to the Pacific plate (Figure 20; Engebretson et al., 1985). Drift was toward the west-northwest, which carried North America some 401 of latitude northward by the Paleocene, before a change in drift trajectories directed the continent as much as 101 back southwestward. For the first time during the Phanerozoic, the northern part of the continent was carried into temperate and then even more northerly latitudes. The Cordilleran orogen has been constructed by accretionary tectonics, which, at least locally, began as far back as the Devonian, and within Canada some 500 km of new continental crust has been added to the western

Figure 18

Plate tectonic setting of North America during the Middle Jurassic.

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Andrew D. Miall and Ronald C. Blakey

Plate tectonic setting of North America during the Early Cretaceous.

Figure 20 The gradual westward drift of North America, relative to Panthalassa, from the Jurassic to the present (Engebretson et al., 1985).

margin of the continent (Price, 1994; Price and Monger, 2003; Ricketts, Chapter 10). Figure 21 provides a series of reconstructions of the plate-tectonic evolution of the western margin of the continent from the Permian to the end of the Cretaceous. These figures are modified from a recent paper by Umhoefer and Blakey (2006) that proposes a compromise on the ‘‘Baja British Columbia’’ hypothesis (see further). From Pennsylvanian through

Phanerozoic Tectonic and Sedimentary Evolution of North America

Figure 21 Evolution of the plate tectonic setting of the western margin of North America from the Permian to the end of the Cretaceous; (A) Permian, (B) Triassic, (C) Early Jurassic (180 Ma), (D) Mid-Jurassic (160 Ma), (E) Late Jurassic (145 Ma), (F) Early Cretaceous (125 Ma), (G) Late Early Cretaceous (105 Ma), (H) Late Cretaceous (85 Ma), (I) Cretaceous-Tertiary boundary (65 Ma).

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(Continued )

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Jurassic, the number of arcs and their polarity with respect to North America is extremely controversial; there is no universal agreement about all of the details of these reconstructions (see Price, 1994; Price and Monger, 2003, for a discussion of the controversies). For example, the faunal data of Smith and Tipper (1991) and the paleomagnetic data of Wynne et al. (1995) suggest that Quesnellia, Stikinia, and Wrangellia all lay considerably further south until the Jurassic than shown here, implying that the structural relationships shown in these diagrams should be modified by substantial right-lateral displacement of post-Jurassic age. This is complicated by the fact that considerable right-lateral displacement is known to have taken place during the Eocene, as discussed below. Partitioning the displacement between the Mesozoic and the Cenozoic is problematic. Collision of the McCloud-Nicola arc against North America took place between the Late Triassic and Early Jurassic (Figures 21A–C), resulting in delamination of the arc and backarc region, and the beginning of a process of obduction and tectonic wedging of the arc eastward over the Precambrian basement. Subduction stepped out to the outboard margin of what was to become the Intermontane belt in Canada (the amalgamated Stikine, Quesnel, and related terranes). The Bridge River terrane of southwest British Columbia, of Permian to MidJurassic age, constitutes the subduction complex and related rocks of the marginal arc (Ricketts, Chapter 10). To the south, the Sierra Nevada arc (along the spine of what was to become California) was established in the latest Triassic, and persisted until the Early Cenozoic (Ingersoll, Chapter 11). Meanwhile, off to the west in Panthalassa, a large terrane now called the Wrangellia superterrane was approaching the North American margin. Wrangellia consists of two large terranes that are thought to have amalgamated in the Mid-Jurassic, within Panthalassa, before suturing to North America (Saleeby, 1983). The Alexander terrane consists of a succession of arc-related rocks resting on a metamorphic basement ranging from Ordovician to Triassic in age. The Wrangell terrane is composed of an arc, overlying plateau basalts, and carbonate sediments of Carboniferous to Triassic age. Both terranes are regarded as far-traveled remnants, of possible Asian origin. Irving et al. (1985) used paleomagnetic evidence to reconstruct the history of amalgamation and accretion, and coined the term ‘‘Baja British Columbia’’ for the configuration and geographic relationship this superterrane would have had until the Mid-Cretaceous (Figures 21D–G). The orientation of the Wrangellia superterrane was oblique relative to the continental margin, such that the southern end of the terrane collided with North America first, in the area of California. This took place in the Mid- to Late Jurassic (Figures 21D and E), causing the Nevadan orogeny, an event which included the emplacement of ophiolite complexes in northern California (Ingersoll, Chapter 11). Collision of Wrangellia with the North American margin was probably not completed until the Mid-Cretaceous, in southern British Columbia (Figures 21F and G). During the Late Jurassic and Early Cretaceous, the Intermontane Superterrane became wedged beneath its outboard margin, and the Cordilleran miogeocline was scraped off its basement and accreted to the advancing front of the Intermontane Superterrane, where it formed the oldest part of the Rocky Mountain foreland fold and thrust belt (Price, 1994). During the collision with the Insular Superterrane, the Intermontane Superterrane (Stikinia, Quesnellia, etc.) was pushed northeastward over the margin of the North American continent, scraping off more of the supracrustal rocks of the Western Canada Sedimentary Basin to produce the rest of the Rocky Mountain foreland fold and thrust belt and the foreland basin (Chapter 9). The accreted terranes were also displaced northwestward, between the Mid-Cretaceous and the Mid-Eocene, producing a set of major right-lateral strike-slip fault systems (Tintina Trench-Northern Rocky Mountain Trench, Shakwak-Denali, Yalakom-Ross Lake, and Fraser RiverStraight Creek faults) that dominate the structural fabric of the Canadian Cordillera. Thus, the Rocky Mountain foreland fold and thrust belt is a transpressional accretionary prism (Price, 1994). The southwestward bend in North America’s path, relative to Panthalassa, that took place at about 60 Ma (Figure 20), explains why transpression became progressively more important from the beginning of the Cenozoic. This culminated in the Eocene with an important phase of northwest-southeast extension and mafic igneous activity, including dyke emplacement, and extrusion of the Kamloops volcanics in central British Columbia. The Western Interior foreland basin was initiated in the Mid-Jurassic as a result of the crustal loading caused by arc collision and terrane amalgamation along the length of the Cordilleran orogen (Miall et al., Chapter 9). Interestingly, the appearance of western sediment sources for this basin seems to have been more or less simultaneous along the length of the basin, at least from Alberta to Utah, despite the diachroneity of the terrane collisions that generated contraction and uplift.

5.2. Sedimentary evolution of the western margin Collisional retroforeland basins are foreland basins developed behind the orogen on the overriding plate at a suture. As Ingersoll (Chapter 11) notes, there do not appear to be any such basins in the Cordillera associated with the terrane collision and amalgamation process. Until the Late Triassic, North America constituted the downgoing plate relative to Panthalassa subduction (Antler and Sonoma orogenies). After arc collapse and

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collision commenced in the Jurassic, basins are difficult to categorize in terms of simple plate models, and may have undergone evolution through several discrete styles. A continental arc was established on the North American margin extending from southern Arizona to northern California late in the Triassic, prior to the collision with Wrangellia (Ingersoll, Chapter 11; Figure 21C) and evolved through several configurations as Wrangellia collided and was displaced northward during the Jurassic–Early Cenozoic (Figures 21D–I). Following the Late Jurassic Nevadan orogeny a major forearc basin, the Great Valley Basin, persisted as a major depository, accumulating a thick succession of sedimentary-gravity-flow deposits (Ingersoll, Chapter 11). Many of the basins within the US and Canadian Cordillera may be described as ‘‘successor basins,’’ that is, basins developed over a deformed orogen. Some are simply residual depressions between elevated blocks, others are formed by subsidence of a previously deformed region of continental crust. For example, uplift of the Cache Creek terrane during the final amalgamation of Quesnellia and Stikinia provided a sediment source for the Bowser Basin (Mid-Jurassic–Lower Cretaceous), which developed over a thermally subsided Stikine basement (Ricketts, Chapter 10). Some of the more significant sedimentary accumulations in the Cordillera are ‘‘overlap assemblages,’’ that is, sediments derived from terranes that had become amalgamated and uplifted, to become a sediment source. The deposition of detritus from one terrane on the eroded surface of the adjacent terrane provides a minimum age for bracketing the time of collision and amalgamation. The Gravina-Nutzotin, Dezadeash, and Gambier basins in central British Columbia help, in this way, to constrain the timing of the collision of the Wrangellia Superterrane with North America (Ricketts, Chapter 10). Several basins, including the Queen Charlotte (Lower–Upper Cretaceous), Georgia (Nanaimo: Mid Upper Cretaceous–Neogene), and Tofino (Paleogene to Recent) basins occupy forearc positions close to the outboard margin of the continent (Ricketts, Chapter 10).

5.3. Sedimentary evolution of the western interior The Western Interior foreland basin is one of the largest and most intensively studied basins in the world (Miall et al., Chapter 9). It has become convenient to regard the Middle to Late Jurassic Nevadan orogeny as the event that initiated the foreland basin, because sediments of this age range in the basin contain the first indication of westerly derived orogenic source. Thickness distribution of the Carmel Formation (Bathonian) in the Colorado plateau area, and the presence of volcanic detritus, suggests derivation from Cordilleran igneous origins and the beginning of a foreland-basin style subsidence pattern (Blakey, Chapter 7). Somewhat later in British Columbia, the collision of Quesnellia with the continental margin caused the uplift from which the Kootenay Formation (latest Jurassic and Early Cretaceous) of Alberta and Montana, was derived (Miall et al., Chapter 9). The Morrison Formation (Mid-Jurassic–Lower Cretaceous) is the first major, widespread clastic unit in the Rocky Mountain States that could be described as a foreland-basin clastic wedge. A particularly important tectonic episode occurred later during the Early Cretaceous, when the entire foreland basin was uplifted and eroded for some 9 Myr, forming a widespread unconformity. A thin but widespread sheet of coarse gravels and sands was then spread across the basin from the rising mountains, which by then extended from northeastern British Columbia to Utah. In Alberta, this gravel, the Cadomin Formation, rests directly on the marine shales of the Fernie Formation (the last pre-orogenic sedimentary unit) above a major unconformity. Geochemical evidence suggests that uplift of the largely oceanic Quesnel terrane and associated magmatic rocks provided a major sediment source for the foreland basin to the east at this time. Lower Cretaceous clastic wedges (Mannville and Dunvegan formations) reflect contraction and uplift caused by the suturing of the Coastal and Insular belts with the continental margin. The world-famous dinosaur country of the Red Deer Valley, Alberta, is located within the fourth of these pulses of clastic sedimentation, the Belly River-Edmonton wedge. The last episode of tectonism reflects the period of transpressive orogenic activity that began in the Late Cretaceous, and led to the deposition of the widespread Paskapoo Formation, of Paleocene age. Deposition continued at least until the Miocene, but only remnants of the youngest part of this final clastic wedge are now preserved, in locations such as the Cypress Hills of southwestern Saskatchewan. These are the only patches left of sediment once carried by river systems as far as the eastern continental margin of Canada. Within a region extending from Montana to northernmost Mexico, the Western Interior was disrupted, beginning in the Late Cretaceous (about 75 Ma), and continuing to the Late Eocene, by basement-involved tectonism, that broke up the foreland basin into a series of relatively small basins and uplifts (the largest, the Powder River Basin, is about 400 km long). This phase of tectonism is quite different from the classic ‘‘thinskinned’’ tectonism of the Sevier and Rocky Mountain (Alberta) fold-thrust belts, and gave rise many years ago to the term Laramide orogeny. The term Laramide has subsequently been used in diverse ways, for example, to refer

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to all Cordilleran orogeny of this age range, a tendency that should be avoided. Various explanations have been advanced for the unusual style of tectonism. Lawton (Chapter 12) discusses the model of ‘‘flat-slab’’ subduction, first advanced by Dickinson and Snyder (1978). The basis of this model is that an area of near-horizontal subduction of a buoyant Pacific (Farallon) plate beneath western North America generated disruption and shortening through the overlying continental crust. Laramide basins are mostly bordered by fault-bounded, basement-cored uplifts. Faults are typically steeply dipping reverse faults showing displacements that may exceed 10 km. Basin fills are almost entirely nonmarine, commonly consisting of coarse fluvial facies. Many of the basins in the center of the Cordillera, from southern Montana to southwest Utah, are classified as ‘‘ponded basins’’ and contain fills that are largely lacustrine in origin.

5.4. Sedimentary evolution of the Arctic margin As noted above, a phase of extension in the Carboniferous triggered the development of Sverdrup Basin in the Canadian Arctic, on the northern margin of the continent. Carbonate and evaporite sedimentation dominated in this basin, together with a belt of marginal clastics, until the Permian, after which, sedimentation became predominantly clastic (Embry and Beauchamp, Chapter 13). Late Paleozoic uplift of the newly formed Pangea (Burgess, Chapter 2), and a northward drift of the continent into higher latitudes, are the probable cause of this major facies change (compare Figures 13, 17, and 18). Sedimentation was brought to an end by the Eurekan orogeny, in the Mid-Cenozoic, during Greenland’s brief life as a separate plate. Arctic paleogeography was profoundly altered by a minor phase of sea-floor spreading in the Early Jurassic. Until then, Canada’s northern Yukon and Arctic Islands region had continued to the north into a region called ‘‘Crockerland’’ (Embry and Beauchamp, Chapter 13). This land area formed a northern margin and sometime sediment source for Sverdrup basin and basins in northern Alaska. Evidence for rifting of Crockerland away from North America is found in fault-bounded basins of Late Jurassic age in the western Canadian Arctic. Sea-floor spreading eventually rotated Crockerland counterclockwise and it collided with, and underthrust northern Alaska, generating the Brooks range in the process (Figure 21F; Embry, 1990). This series of events generated a new extensional continental margin for Canada, where the northwest margin of the Arctic Islands is now located (Embry and Beauchamp, Chapter 13). To the southwest, this passes into the Beaufort-Mackenzie Basin (Dixon et al., Chapter 16). The east side of this basin is comparable in structure and stratigraphy to the extensional Arctic Islands margin, whereas the western part of the basin was affected by thin-skinned tectonism between the latest Cretaceous and the Late Miocene. The basin there has the character of a foreland basin adjacent to the Brooks Range fold-thrust belt. Active sedimentation continues in the Beaufort-Mackenzie Basin. It is currently the depository for the Mackenzie River, one of the major rivers draining the North American interior.

5.5. Sedimentary evolution of the Atlantic and Gulf margins The Atlantic margin was initiated as a series of rift basins, including the well-known Newark rifts, of Triassic age (Figure 17). Sedimentation of coarse clastics along basin margins passed, in many cases, into evaporites in basin centers (Miall et al., Chapter 14). Evaporite sedimentation persisted into the Jurassic throughout the central Atlantic, from Newfoundland and Nova Scotia (Argo Salt) to the Gulf of Mexico (Louann Salt). At this time, the incipient Atlantic-Gulf ocean lay between about 151N and 251N latitude, and had minimal connections to the global oceanic circulation system (Figure 18). Rifting extended into the northern Atlantic margins and the Baffin Bay-Davis Strait area in the Early Cretaceous. Sedimentation there was predominantly clastic. Evaporite sedimentation did not occur in these more northerly basins, which were located in cooler climates. During most of the Jurassic and into the earliest Cretaceous, carbonate sedimentation extended along much of the Atlantic margin (Miall et al., Chapter 14) and into the Gulf Coast basin (Galloway, Chapter 15), where such classic carbonate units as the Smackover Formation were deposited. However, as North America drifted slowly northwestward (Figure 20), this style of sedimentation came to an end. It ended by Neocomian time on the Scotian Shelf and Georges Bank, and by Cenomanian time on the Blake Plateau, to be replaced mainly by shallow-marine clastics (Jansa, 1981). On the Bahamas platform and part of the Florida margin, carbonate sedimentation continues to the present day. Loading and displacement of the basal salt layers had profound consequences for the evolution of the Gulf Coast basin, as discussed in detail by Galloway (Chapter 15). This basin was where modern North American petroleum geology found its feet, in the early part of the twentieth century. Prospecting for evaporite diapirs using primitive gravity and seismic devices was one of the first applications of scientific methods, anywhere, to the exploration for petroleum. The Gulf Coast is the depository for several of the major river systems draining the interior of North America, including the Mississippi-Missouri and the Rio Grande. These have remained in essentially the same

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positions since the Early Cenozoic (Fisher and McGowen, 1967). Sedimentation has extended the continental margin of the Gulf by more than 300 km since the Early Cretaceous. Sedimentary pulses may be correlated to major tectonic episodes within the continental interior, especially episodes of deformation and uplift of the Cordilleran orogen (Galloway, Chapter 15). Laramide deformation commenced in the Campanian (Lawton, Chapter 12), but the main phase of deformation, as indicated by coarse clastic supply into Laramide basins, extended from Maastrichtian to Late Eocene, roughly contemporaneous with the Wilcox depositional episode in the Gulf Coast (Chapter 15). Uplift and erosional truncation of Laramide structures indicates the end of this tectonic episode by Late Eocene time, with subsequent deposition of the Early Oligocene Bishop Conglomerate in the Uinta Basin contemporaneous with the beginning of the Frio-Vicksburg depositional episode in the Gulf Coast Basin (Chapter 15). At about 40 Ma, subduction of the Panthalassa Ocean brought the sea-floor spreading ridge between the Pacific and Farallon plates against the California continental margin. Relative motions between the North American and Farallon plates were oblique, and subduction of the spreading ridge led to the development of the right-lateral San Andreas transform fault (Atwater and Molnar, 1973). Changes in intraplate stress across the western part of the continent that were triggered by this event are now thought by most workers to be the main cause of the change from a contractional to an extensional regime within the US Cordillera. A major phase of crustal extension took place between the Oligocene and the Miocene, leading in places to a doubling of the width of some regions of continental crust. The resulting tectonic province, the Basin and Range, extends from southern British Columbia to northern Mexico (Oldow et al., 1989). Sedimentary basins in this province are characteristically of graben or half-graben style, bounded by steep listric faults that typically flatten out at midcrustal levels (Hamilton, 1987). Basin fills are entirely nonmarine. Many small basins bounded by strike-slip faults are located within the San Andreas transform system (Nilsen and Sylvester, 1995), including nonmarine basins onland (e.g., Crowell and Link, 1982), and offshore basins that are largely filled with turbidites and other deep-marine clastics (e.g., Howell et al., 1980).

6. Late Cenozoic Modifications Continental glaciation during the Late Cenozoic spread a thick blanket of moraine and outwash deposits across the northern interior, as far south as Illinois. Glaciation disrupted continental drainage patterns. A major east-flowing river system is thought to have crossed the continental interior until the Late Cenozoic, transporting detritus from the northern Cordillera to the continental margin of Baffin Bay (McMillan, 1973; Duk-Rodkin and Hughes, 1994). This river system was dammed and diverted by advancing continental ice. Outflow found a course to the northwest, and this is probably the origin of the present-day Mackenzie River system. Post-glacial reworking of the glacial blanket has triggered major slumping and gully erosion of the Baffin Bay margins (Chough and Hesse, 1976). Meanwhile, tidewater glaciers continue to shed vast quantities of clastic debris into the Gulf of Alaska (Eyles et al., 1991). Melting of the continental ice caps generated enormous proglacial lakes (including Bonneville, Missoula, Agassiz, Iroquois, and other super-Great Lakes), the draining of which generated major meltwater channels and carried significant volumes of sediment out to the Pacific, Arctic, and Atlantic margins. During part of the meltback phase, the Great Lakes drained southward through the Mississippi system (Rittenour et al., 2007), and was responsible for depositing a giant submarine fan system on the floor of the Gulf of Mexico (Feeley et al., 1990; Weimer, 1990). Following the breakup of Rodinia, the newly separated continent, Laurentia, was subjected to profound erosion and peneplanation. The Grenville orogeny created a major mountain belt in eastern North America about a billion years ago, but by the time of the major Early Paleozoic transgressions, this had almost completely disappeared. Some evidence suggests that the thick Neoproterozoic clastic wedges on the western and northern margins of Laurentia (Figure 14) may have been derived in part by peneplanation of the Grenville mountains (Rainbird et al., 1997; see also Chapter 17). Tracing the present erosional edge of the Phanerozoic cover around the margins of the Canadian Shield, a geologist cannot help but be struck by the lack of topographic relief on the unconformity surface. Basal Paleozoic strata contain a few sandstone layers and scattered boulders, but in many places, a few meters above the unconformity, the section passes up into platform carbonates, evidence for the development of giant, detritus-free, interior platform seas of a magnitude nowhere replicated on the present-day Earth’s surface. Reuter and Watts (2004) described a buried paleochannel on the Precambrian surface in Ohio, one of the few documented exceptions to this general rule. The development of this enormous peneplain suggests that the continent remained largely undisturbed by tectonism for at least 200 Myr, from the end of the Grenville orogeny until the beginning of rifting at about

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800 Ma. The continental interior continued to undergo subaerial erosion until the first Paleozoic transgressions, some 300 Myr later. As noted above, dynamic topographic processes generated broad westward or eastward tilts to the continent at different times during the Late Precambrian and Phanerozoic, which controlled fluvial erosion, sediment transport and sediment delivery on a continental scale, and provided for the removal of enormous volumes of rock in the construction of the post-Precambrian peneplain and the major Phanerozoic sequence boundaries (see also Chapter 17). Large areas of North America may be at the beginning of another long period of peneplanation. The Appalachian orogen, in the east, is already exposed to its roots. In the west, some 3 km of the sediment-fill have been removed from the Alberta Basin, following post-orogenic uplift (Beaumont, 1981). The Mid-Cenozoic extension of the southwest Cordillera to form the Basin and Range region initiated a significant reduction in relief in that area. By contrast, although there is an absence of obvious targets for terrane collision in the Pacific Ocean, continued subduction of the Gorda and Pacific plates beneath the western margin, from northern California to Alaska is likely to maintain active tectonic uplift in the northwest. Uplift of the Colorado Plateau is an exception to the general trend, and possible future activity of the Yellowstone hot spot may complicate this simple picture, but on the very long term, the rest of the continent seems likely to have entered another hundreds-of-millions-of-years long phase of erosional reduction.

ACKNOWLEDGMENTS This chapter has been reviewed by all the contributing chapter authors or senior co-authors of this book, who have pointed out errors and suggested a number of useful points of clarification.

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Srivastava, S. P., Falconer, R. K. H., and MacLean, B., 1981, Labrador Sea, Davis Strait, Baffin Bay; geology and geophysics, a review, in Fergusson, A. J. and Kerr, J. W. eds., Geology of the North Atlantic Borderlands, Canadian Society of Petroleum Geologists, Calgary, AB, Memoir, v. 7, pp. 333–398. Stampfli, G. M., von Raumer, J. F., and Borel, G. D., 2002, Paleozoic evolution of pre-Variscan terranes: from Gondwana to the Variscan collision, in Martinez Catalan, J. R., Hatcher, R. D., Arenas, R., and Diaz Garcia, F. eds., Variscan-Appalachian dynamics: the building of the Late Paleozoic basement, Geological Society of America (Special Paper 364), Boulder, CO, pp. 263–280. Surlyk, F., Clemmensen, L. B., and Larsen, H. C., 1981, Post-Paleozoic evolution of the East Greenland continental margin, in Kerr, J. W. and Fergusson, A. J. eds., Geology of the North Atlantic Borderlands, Canadian Society of Petroleum Geologists, Calgary, AB, Memoir, v. 7, pp. 611–646. Tankard, A. J., and Welsink, H. J., 1987, Extensional tectonics and stratigraphy of Hibernia oil field, Grand Banks, Newfoundland. American Association of Petroleum Geologists Bulletin, v. 71, pp. 1210–1232. Thomas, W. A., 2006, Tectonic inheritance at a continental margin. GAS Today, v. 16(2), pp. 4–11. Trettin, H. P., 1989, The Arctic Islands, in Bally, A. W. and Palmer, A. R. eds., The geology of North America: an overview, The Geology of North America, Geological Society of America, Boulder, CO, v. A, pp. 349–370. Umhoefer, P. J., and Blakey, R. C., 2006, Moderate (1600 km) northward translation of Baja British Columbia from southern California: An attempt at reconciliation of paleomagnetism and geology, in Haggart, J. W., Enkin, R. J. and Monger, J. W. H., eds., Paleogeography of the North American Cordillera: Evidence for and against Large-Scale Displacements. Geological Association of Canada, St. John’s, Nfld. (Special Paper 46), pp. 305–327. Van Staal, C. R., Dewey, J. F., Miocaill, C. M., and McKerrow, W. S., 1998, The Cambrian-Silurian tectonic evolution of the northern Appalachians and British Caledonides: history of a complex west and southwest Pacific-type segment of Iapetus, in Blundell, D. J. and Scott, A. C. eds., Lyell: the past is the key to the present, Geological Society of London (Special Publication), London, UK, v. 143, pp. 199–242. Weimer, P., 1990, Sequence stratigraphy, facies geometries, and depositional history of the Mississippi fan, Gulf of Mexico. American Association of Petroleum Geologists Bulletin, v. 74, pp. 425–453. Williams, H., 1978, Tectonic lithofacies map of the Appalachian orogen, Memorial University of Newfoundland Map 1, scale 1:1,000,000. Worrall, D. M., and Snelson, S., 1989, Evolution of the northern Gulf of Mexico, with emphasis on Cenozoic growth faulting and the role of salt, in Bally, A. W. and Palmer, A. R. eds., The geology of North America: an overview, The Geology of North America, Geological Society of America, Boulder, CO, v. A, pp. 97–138. Worsley, T. R., Nance, D., and Moody, J. B., 1986, Tectonic cycles and the history of the earth’s biogeochemical and paleoceanographic record. Paleoceanography, v. 1, pp. 233–263. Wynne, P. J., Irving, E., Maxson, J. A., and Kleinsphehn, K. L., 1995, Paleomagnetism of the Upper Cretaceous strata of Mount tallow: evidence for 3000 km of northward displacement of the eastern Coast Belt, British Columbia. Journal of Geophysical research, v. 100, pp. 6073–6091. Yorath, C. J., 1991, Upper Jurassic to Paleogene assemblages; Chapter 9, in Gabrielse, H., and Yorath, C. J. eds., Geology of the Cordilleran Orogen in Canada, Geology of Canada, Geological Survey of Canada, Ottawa, N. 4, pp. 329–371. Ziegler, P. A., 1988, Evolution of the Arctic-North Atlantic and the Western Tethys. American Association of Petroleum Geologists, Memoir, v. 43, pp. 198.

CHAPTER 2

Phanerozoic Evolution of the Sedimentary Cover of the North American Craton Peter M. Burgess

Contents 1. Introduction 2. Deﬁnition of a Craton 3. Tectonic Elements of the North American Craton 3.1. The Canadian shield 3.2. The Cratonic platform 3.3. Intracratonic basins 3.4. Cratonic margins 4. Controls on Evolution of the Cratonic Cover 4.1. Eustasy 4.2. Extension and thermal re-equilibration 4.3. Intraplate stress 4.4. Dynamic topography related to subducting slabs 4.5. Dynamic topography related to mantle insulation and supercontinent cycles 4.6. Mantle downwelling: ‘‘cold spots’’ 4.7. Magmatic controls 5. Phanerozoic Evolution of the Cratonic Platform Cover 5.1. The Sauk sequence (Late Precambrian to Early Ordovician) 5.2. The Tippecanoe sequence (Middle Ordovician to Early Devonian) 5.3. Kaskaskia sequence (mid-Early Devonian to Late Mississippian) 5.4. Absaroka sequence (Late Mississippian to Early Jurassic) 5.5. Zuni sequence (Middle Jurassic to Early Paleocene) 5.6. Tejas sequence (Late Paleocene to present) 6. The North American Intracratonic Basins 6.1. The Michigan Basin 6.2. The Illinois Basin 6.3. The Williston Basin 6.4. The Hudson Bay Basin 7. Summary Acknowledgments References

32 32 32 33 34 34 35 36 36 38 38 39 40 40 41 43 43 44 46 46 49 50 51 51 53 56 58 61 61 61

Abstract Although the term craton is often taken as synonomous with tectonic quiescence, the North American craton is not simply an unchanging, stable platform accumulating strata and inﬂuenced only by changes in global sea-level. Rather, viewed on a timescale of tens to hundreds of millions of years at least, it is a dynamic tectonic environment inﬂuenced by various plate tectonic, mantle, denudational and depositional processes. The Sloss cratonic sequences record the history of this dynamic tectonic environment, in the form of episodes of transgression, regression and erosion and non-deposition, generated on a timescale of tens of millions of years. These sequences occur across the craton, on areas of platform, as well as in the four main intracratonic basins, yet their origins remain relatively poorly understood. Long-term eustatic oscillations must certainly have contributed to development of the transgressive and regressive sequence elements, but basic observations of tilted strata and angular sequence-bounding unconformities show eustasy cannot have been the only responsible mechanism. Variations in dynamic topography generated by subducting lithospheric slabs, and by thermal insulation of mantle beneath supercontinents, can explain much of the large-scale sequence architecture but more detailed plate tectonic reconstructions and associated mantle convection models are necessary to further test and Sedimentary Basins of the World, Volume 5 ISSN 1874-5997, DOI 10.1016/S1874-5997(08)00002-6

r 2008 Elsevier B.V. All rights reserved.

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develop these explanations. Intraplate stress also seems likely to have played a large role in generating the cratonic sequences by reactivating pre-existing structures and driving subsidence and uplift. Variations in intraplate stress through time can be related, to some degree at least, to tectonic events occurring on the cratonic margins and on other adjacent plate margins. Given present available evidence and theory, the North American intracratonic basins seem most likely to be due to a combination of mantle downwelling and focused intraplate stress variations, in some cases with an element of long-wavelength tilting due to subduction-induced dynamic topography, and in some cases with an initial trigger by lithospheric stretching. Although taken together all these mechanisms provide a plausible explanation for the development of the North American cratonic sequences, they are certainly not deﬁnitive, conclusive explanations. Much work remains to be done to test, and to conﬁrm or refute these ideas.

1. Introduction The North American craton extends north to south for more than 5,500 km from mid continental USA to the Arctic Islands, and for more than 4,700 km west to east, from western Canada to eastern Greenland. Given the geographical extent of this geological feature, the diversity of strata preserved on it, and the volume of data produced by more than one hundred years of study, it is perhaps surprising that the Phanerozoic history of the craton is often viewed as, by deﬁnition, ‘‘tectonically stable.’’ The history recorded in the cratonic Phanerozoic strata suggests a rather more complex evolution. The craton was sometimes tectonically quiescent, sometimes ﬂooded almost entirely to produce vast shallow seas like nothing existing on Earth today, sometimes being uplifted, warped and tilted with material being stripped off by erosion at geologically rapid rates. This chapter describes some of the diversity of structure and strata preserved on the craton, and provides an interpretation of its history. It lends support to a view (e.g. Sloss and Speed, 1974) of the North American craton as a dynamic geological entity with a varied and sometimes spectacular history that can enhance our understanding of Earth surface and mantle processes.

2. Deﬁnition of a Craton In simple terms a craton is a stable, strong, unyielding area of lithosphere. The term is difﬁcult to deﬁne in more detail (Leighton, 1990) because cratonic areas usually have a history of intense Precambrian deformation, followed by relative tectonic quiescence during Phanerozoic time. This history reﬂects the ongoing tectonic process of continental accretion; parts of the craton now safely ensconced within the continent were, in Proterozoic time, areas of tectonic activity on the craton margin. Currently active continental margins may in the future form part of a stable cratonic basement when the next phase of orogenesis forms a new accreting, active margin. Conversely, areas currently within the craton may become zones of renewed plate-margin tectonic activity due to rifting and oceanic spreading. The term craton was ﬁrst used in the mid-20th century, during development of ideas about geosynclines, to contrast the relatively young, active continental margins with the more ancient, apparently less active continental interiors. Problems arose when geologists attempted to identify features marking the boundary between craton and geosynclines and realized that in terms of stratal thickness and characteristics, the difference is difﬁcult to deﬁne (Sloss, 1988a). Developing understanding of plate tectonics resolved this problem by allowing a distinction based on the underlying subsidence mechanism. Areas formerly termed geosynclines were redeﬁned as rifted margins undergoing thermal subsidence, or active-margin basins formed by a variety of subsidence mechanisms including rifting and lithospheric loading. The craton was then assumed to be the area of old, strong lithosphere beyond the reach of these marginal tectonic mechanisms (Sloss, 1988a). For the purposes of this chapter, and bearing in mind that the above is a time-limited deﬁnition, the craton in North America is considered to be the area underlain by Precambrian basement that has not been subject to platemargin processes during the Phanerozoic (Figure 1). Note that this does not mean that the craton has not been subjected to signiﬁcant tectonic uplift, subsidence and deformation, but simply that the tectonic processes affecting cratonic interiors are epeirogenic, and so do not result directly as a consequence of active plate-margin processes.

3. Tectonic Elements of the North American Craton Subdividing the craton into different elements serves to illustrate the different tectonic mechanisms operating within the craton, and to illustrate progressive cratonic evolution.

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Figure 1 Map of North American cratonic tectonic elements. Modi¢ed from Bally (1980). Tectonic element names taken from Stott and Aitken (1993) and Sloss (1988b).

3.1. The Canadian shield A shield is part of a craton, and is deﬁned by the American Geological Institute Glossary of Geology as a large region of exposed basement rocks, commonly with a very gently convex surface, surrounded by sedimentcovered platform. In North America the shield covers a large part of eastern and central Canada, and also covers most of Greenland, although currently buried beneath the Greenland ice cap (Figure 1). Shield rocks consist of

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predominantly granitic gneises, and greenstone belts composed of metasedimentary and metavolcanic rocks, and can be subdivided into seven provinces with different geological histories, based on spatial distribution, crosscutting relationships and isotopic ages (Hoffman, 1989). Did the shield persist in its present state, subaerially exposed and without any sedimentary cover, throughout the Phanerozoic? This would imply absence of tectonic uplift or subsidence in a stable, quiescent setting. However, there is evidence that much of the shield has been covered with platform strata for certain times during the Phanerozoic, and has undergone uplift and erosion at other times. Sloss (1988b) and Sloss (1963) point out that Middle and Upper Ordovician distal marine strata are in many places terminated by erosional truncation beneath overstepping Devonian strata, strongly suggesting that the Ordovician and Devonian strata covered a signiﬁcantly greater portion of the shield before being removed by erosion. Other examples of similar stratal terminations are recorded by Bunker et al. (1988) and Collinson et al. (1988). Limestone xenoliths found in kimberlites on the Canadian shield are Middle Devonian in age, and appear to record deposition of more than 750 m of Devonian and younger strata, preserved on the craton until Middle Jurassic time, and removed by erosion prior to the Late Cretaceous epoch (Cookenboo et al., 1998). Nd isotope data show that the shield was a dominant source of cratonic sediment during Early Paleozoic time (Patchett et al., 1999), concordant with observations from Sloss (1988b), but much of this sediment probably originated from further north than the present southern limit of the shield (Collinson et al., 1988). Patchett et al. (2004) also used Nd isotope data from strata from the Sverdrup basin of the Canadian Arctic Islands to show that much of the shield was probably covered by Ordovician to Middle Devonian carbonate units, the northerly derived Upper Devonian siliciclastic sedimentary rocks probably covered about one-half of the shield in its western and northern portions and that this cover was progressively removed through Mesozoic time. Taken together, this evidence clearly indicates epeirogenic activity on the shield throughout the Phanerozoic leading to cycles of ﬂooding, re-emergence and tilting.

3.2. The Cratonic platform The term platform is deﬁned in the American Geological Institute’s Glossary of Geology as a part of a continent covered by ﬂat-lying or gently tilted sedimentary rocks, underlain by a complex of rocks that were consolidated during earlier deformations. Platform areas can be considered sedimentary basins in the sense that they are areas of sediment accumulation, much of which is now at depths greater than when deposited (Aitken, 1993). However, they are not basins in the sense of an area undergoing differential subsidence relative to the surrounding area of stable basement. This latter deﬁnition applies to intracratonic basins, but not to cratonic platforms. Cratonic basins within the platform show signiﬁcantly greater thicknesses of preserved strata than do the surrounding platform areas e.g. W4.5 km of strata in the Michigan basin, compared to B1 km of strata on the surrounding platform area (Figure 2). Subsidence mechanisms are not well understood in either case (see Section 4), but there is a clear distinction in terms of thickness of preserved strata. A signiﬁcant feature of the cratonic platform is the presence of epeirogenic arches. An epeirogenic arch is an intraplatform high that subsides less rapidly than surrounding platform areas, leading to formation of relatively thin strata, or is uplifted, leading to erosion and local unconformity. Arches are signiﬁcant because they separate platform areas and appear to have acted as important elements in the paleogeography of the craton. They are also important because they provide a clear indication of epeirogenic tectonic activity within the craton throughout Phanerozoic time. Although arches may subdivide the platform at particular points in time, and areas enclosed by arches are sometimes referred to as intracratonic basins, such areas are still best considered as platform rather than intracratonic basin because of the obvious differences in total preserved sediment thickness (Figure 2).

3.3. Intracratonic basins Intracratonic basins are areas on the craton, at some distance from the craton margin, undergoing differential subsidence relative to the surrounding area of cratonic basement. They are thus distinguished from platform areas by signiﬁcantly greater thicknesses of preserved strata e.g. W4.5 km of strata in the Michigan basin, compared to B1 km of strata on the surrounding platform area. Four North American intracratonic basins are identiﬁed on this basis, namely the Michigan, Illinois, Williston and Hudson Bay basins (Figures 1 and 2). Evolution of the Michigan, Illinois and Williston basins as discrete basins began in Late Cambrian time, suggesting perhaps that they are in some way associated with the break up of a Late Precambrian supercontinent. Initiation of the Hudson Bay basin occurred in Ordovician time, so a link with supercontinent break up is more difﬁcult to deﬁne. All four basins were receiving sediment throughout much of Early Paleozoic time, and were inﬂuenced to varying degrees by marginal and craton-wide tectonic events.

Phanerozoic Evolution of the Sedimentary Cover

Figure 2

35

Map of stratal thickness on the North American craton. Modi¢ed from Sanford (1987).

3.4. Cratonic margins Tectonic mechanisms and events affecting the cratonic margins are varied in their duration and spatial extent, so that there is commonly no clear distinction between marginal and epeirogenic processes. For example, the Colorado plateau is a largely intact cratonic block surrounded by an area of deformation related to

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Figure 3 Chronostratigraphic (Wheeler) diagram (modi¢ed from Catacosinos et al., 1990) showing the six cratonic depositional sequences de¢ned by Sloss (1963). Absolute ages are those used in Sloss (1988b).

Sevier-Laramide compressional and late Cenozoic basin and Range extensional tectonics (Dumitru et al., 1994). Despite its apparent cratonic nature, the plateau area has been subject to active-margin processes at least since the Jurassic period (Allen et al., 2000) (Blakey, Chapter 7; Miall et al., Chapter 9). Similarly, cratonic areas in eastern North America were subjected to margin processes such as ﬂexural loading and slab-related effects from Ordovician time onwards during the Taconic, Acadian and Alleghanian orogenies (Figure 3) (Beaumont et al., 1987; Coakley and Gurnis, 1995; see also Lavoie, Chapter 4; Ettensohn, Chapter 4). However, there is an identiﬁable difference in the magnitude at which these various tectonic processes operated. Stratal thicknesses increase into the cratonic margins (Figure 4) as a result of higher rates of subsidence on the margins. The various cratonic-margin basins are described throughout the rest of this book; in this chapter the inﬂuence of certain marginal tectonic episodes on cratonic strata will be described.

4. Controls on Evolution of the Cratonic Cover Although studied for over one hundred years, the mechanisms responsible for generating the North American cratonic sequences, both on across the platform, and within the four intracratonic basins, are still relatively poorly understood. This is especially true for the cratonic platform where it is difﬁcult even to explain the basic observation of more than 1 km of Phanerozoic, predominantly Paleozoic strata that covers the platform; the mechanisms responsible for generating this long-term cratonic platform subsidence remains poorly understood (Burgess and Gurnis, 1995). In the broadest sense, there are two important candidate processes; eustatic changes in sea-level, and tectonic uplift and subsidence. However, this is a simplistic view because longterm eustatic changes are themselves driven by tectonic processes, namely changes in oceanic spreading and subduction rates, and cratonic tectonic uplift and subsidence may occur by a number of different processes. Also, both eustatic and tectonic processes can themselves trigger a cascade of depositional, denudation and isostatic processes, all of which can be difﬁcult to decipher. A range of these tectonic processes is described here, along with some evidence for and against their operation.

4.1. Eustasy Deﬁnition of the cratonic sequences of North America by Sloss (1963) provided the foundation for much of the sequence stratigraphic model developed since (e.g. Vail et al., 1977). Vail and co-workers chose to emphasize eustasy as the primary mechanism responsible for the relative sea-level changes necessary to explain cratonic

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37

Figure 4 Diagrammatic cross-sections of Sauk,Tippecanoe and Kaskaskia strata illustrating thinning of the lower Paleozoic passive margin strata onto the craton, development of arch and basin geometries on the craton and the often composite nature of the megasequence-bounding unconformities. Note that the datum in the section is the base of the Absaroka sequence. Redrawn from Bally (1989).

Figure 5 Cross-section through the southern mid-continent region of the USA, showing the development of the Sloss sequences in the area, the in£uence of the Nemaha Uplift on stratal patterns, and the thickening-to-the-west wedge of Cretaceous strata developed over slab-related dynamic topography. Modi¢ed from Bunker et al. (1988).

sequence development. Indeed, much of the Paleozoic eustatic curve in Vail et al. (1977) is derived from North American cratonic stratal patterns. However, there is a very basic observation that demonstrates that eustasy was not the only contributor to the relative sea-level changes recorded by cratonic sequences. Long-term eustatic oscillations certainly must have contributed to development of the transgressive and regressive sequence elements, but long-wavelength postdepositional tilting of the cratonic strata and angular sequence-bounding unconformities, both ubiquitous features of North American cratonic strata (e.g. Figures 4 and 5), obviously require a tectonic mechanism, and cannot be explained by eustatic change alone. Applying such simple reasoning to North American cratonic stratal patterns often provides a reasonable indication of the degree of tectonic and eustatic inﬂuence on relative sea-level

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ﬂuctuations. Note that in this case, a tectonic component in the relative sea-level change usually refers to local or regional tectonics causing surface uplift and subsidence. On a larger scale, ridge spreading rate inﬂuences rate of subduction, which, through dynamic topography, inﬂuences ocean volume and hence eustasy. Therefore, on large spatial scales, and at timescales of more than a few million years, distinctions between tectonic and eustatic forcing may be misleading — they are different aspects of the same process (Gurnis, 1990).

4.2. Extension and thermal re-equilibration Stretching of the lithosphere by extensional stress creates extensional sedimentary basins, formed by a combination of initial subsidence due to active normal faulting, followed by subsidence due to post-rift thermal re-equilibration (McKenzie, 1978). This model has been widely applied and used to explain various intracratonic basins, including many of the North American examples (Haxby et al., 1976; Sleep et al., 1980). The Illinois basin is underlain by a Precambrian rift system (Braille et al., 1982), so there is a clear mechanism for subsequent Early Paleozoic thermal subsidence. Later subsidence in the Illinois basin cannot so easily be explained by extensional stress because it occurred beyond the time required for thermal re-equilibration of the lithosphere– asthenosphere boundary (Quinlan, 1987). The Michigan basin is also underlain by a rift system, but it is similarly too old to have inﬂuenced Early Paleozoic subsidence without further extensional episodes, for which there is no direct evidence (Nunn et al., 1984). There is only indirect evidence for existence of a rift system beneath the Williston basin (e.g. Kent, 1987) and the Hudson Bay basin (Quinlan, 1987; Roksandic, 1987). The conclusion, assuming that rift structures are truly absent, is that extensional tectonics may be important in initiating some intracratonic basins, but cannot be invoked alone as a general mechanism for basin formation or for long-term intracratonic basin subsidence (Quinlan, 1987).

4.3. Intraplate stress Stress due to various tectonic events and applied at plate margins can be transmitted laterally through the plates because the lithosphere has some rigidity. Such stress is known as in-plane intraplate stress (Cloetingh, 1986). Cloetingh (1986) pointed out that a change in stress of more than 1 Kbar can produce 50 m of vertical displacement of the lithosphere. This occurs by amplifying pre-existing curvature of the lithosphere, transferring in-plane stress to vertical movements. As well as causing regional uplift and subsidence, stress transmitted through the lithosphere may also cause reactivation of previously existing structures (e.g. Braun and Shaw, 2001; Marshak et al., 2003). Reactivation of faults may regenerate old basement highs and partly invert basins, or cause new periods of extension and subsidence. Consequently, intraplate stress variations may be responsible for onset or cessation of deposition in the cratonic interior, contributing to formation of cratonic sequences (Braun and Shaw, 2001; Burke et al., 2003). Changes in stress of this magnitude are likely to occur during periods of major plate reorganization, causing uplift and subsidence on cratonic margins and in cratonic interiors via intraplate stress and reactivation of old structures (Quinlan, 1987; Ziegler, 1988; Marshak and Paulsen, 1996; Marshak et al., 2003). This would suggest a genetic link between tectonic activity at the cratonic margin and cratonic sequence development, as suggested by various authors (e.g. Sloss and Speed, 1974; Burke et al., 2003). One example of this process is the formation of the Michigan and Williston basins as discrete entities in Late Silurian time, during closure of the Iapetus ocean (Leighton and Kolata, 1990). Another example is the cratonic tectonism associated with the sub-Absaroka unconformity (see Section 5.4), generated by collisional plate convergence in the Marathon and Ouachita orogenesis (Sloss, 1988b). Large-scale tectonic features of the craton such as the Transcontinental Arch and the Nemaha Uplift (Figures 1, 4 and 5) are also evidence of intraplate stress, particularly since the arch had a complex history of uplift and subsidence throughout the Paleozoic that may be linked to inherited basement structure reactivated by variations in intraplate stress (Marshak et al., 2003). Intraplate stress as a mechanism for unconformity generation has been criticized because synchronous activity across the entire craton may be difﬁcult to explain by intraplate stress, since it would require very large stress ﬁelds to develop at a point in time across the whole craton. However, synchronous in the context of cratonic sequences may only mean within a few million years, so this may be less of a weakness in the intraplate stress model than it ﬁrst appears. Also, examples of this kind of synchronous effect have been documented. On the African plate, obduction of ophiolites onto the northeastern margin of Afro-Arabia appears to have generated an unconformity related to intraplate deformation, during a period of 2 Myr in the Senonian, across more than 20 million km2, from Oman to the Atlas, and from Kenya to Equitorial Guinea (Guiraud and Bosworth, 1997; Burke et al., 2003). Given this, application of the intraplate stress model as a mechanism for unconformity generation in the Phanerozoic for North America does not seem unreasonable.

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39

4.4. Dynamic topography related to subducting slabs Subducting slabs are cold and therefore dense relative to the surrounding mantle, forming a positive mass anomaly within the mantle. This positive mass hangs in the mantle and exerts a force on the overlying lithosphere. The force is transferred via viscous mantle ﬂow and acts on the base of the lithosphere, dragging it down to cause a dynamic topographic low. Such topography is considered to be dynamic because it is generated by mantle buoyancy forces that are constantly changing, due to convection within the mantle, but slowly enough that the dynamic topography can be considered in equilibrium with the evolving forces at any instant in time. A subducting slab will therefore produce a region of depressed lithosphere, extending as much as 2,000 km into the craton, from a maximum near the active continental margin (e.g. Mitrovica et al., 1989; Gurnis, 1993) (Figure 6). Such a dynamic topographic low will develop by subsidence as the subducting slab penetrates the mantle. The amplitude of the topography will depend on the degree of slab penetration into the mantle, and on the temperature of the slab; older oceanic crust will be colder and denser than younger crust, representing a greater mass anomaly in the mantle, and generating a higher amplitude of dynamic topography. Consequently, dynamic topography will peak at the point of maximum slab penetration and thermal age (Burgess and Moresi, 1999). The dynamic topography will then gradually be reduced, causing uplift, as the slab’s thermal age decreases during ﬁnal stages of ocean subduction, and/or the slab is detached and descends into the mantle (Gurnis, 1993; Burgess and Moresi, 1999). Since North America had several episodes of subduction at its margins during Phanerozoic time, a contribution to the tectonic evolution of the craton from slab-related dynamic topography is to be expected. Burgess et al. (1997) studied this effect, with particular emphasis on two episodes: subduction of Iapetus oceanic lithosphere during Early Paleozoic time, and subduction of Paciﬁc oceanic lithosphere during the Mesozoic and Cenozoic Cordilleran orogeny. We found that the Early Paleozoic subduction episode probably contributed to development of Lower Paleozoic strata in eastern North America, as also described by Coakley and Gurnis (1995) but a more detailed subduction history is necessary to elaborate further. In contrast, Mesozoic and Cenozoic slab evolution is better constrained. In this case combined mantle and stratigraphic modelling highlights the likely role of slab-related dynamic topography in the development of Cordilleran and cratonic stratal patterns in western North America (e.g. Burgess et al., 1997; Burgess and Moresi, 1999; Cross, 1986; Cross and Pilger, 1978; Pang and Nummedal, 1995; Lie and Nummedal, 2004). For example, development of the Cretaceous Interior Seaway and distribution of Cretaceous strata across a large area of the craton was coeval with penetration of the near-horizontal Farallon slab beneath southwestern North America in Late Cretaceous and Early Cenozoic time (see Miall et al., Chapter 9). One explanation for the wide extent of Cretaceous strata suggests that a dynamic topographic low produced by this subduction created accommodation across much of the craton, and combined at the cratonic margins with other tectonic mechanisms such as elastic loading by orogenic wedges. Accumulation of Cretaceous and Tertiary strata in this dynamic topographic low was

Figure 6

Mechanisms generating dynamic topography driving vertical motions of the cratonic interior.

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followed by surface uplift during reversal of the topography as the slab thermal age decreased, slab dip increased and ﬁnally the slab detached. This is a possible explanation for why much of these near-horizontal strata are now elevated signiﬁcantly above sea-level and being eroded (Burgess et al., 1997). A weakness with the slab-related dynamic topography model has been the lack of obvious dynamic topography on modern convergent margins, such as in South East Asia. In this area various mantle ﬂow models predict 1–2 km of dynamic topography yet the maximum amplitude of dynamic topography in South East Asia has been calculated as no more than 300 m (Wheeler and White, 2000). However, Husson (2006) demonstrates that a signiﬁcant fraction of the topographic variations observed above the Scotia, Mariana and Hellenic subduction systems is consistent with simple theoretical models of dynamic stresses induced by subduction. From this, Husson (2006) suggests that slab-related dynamic topography is difﬁcult to detect in modern tectonic systems only because isostatic processes tend to mask it. Thus, interpretations of the relative importance of slab-related dynamic topography to North American cratonic stratal patterns is open to debate, but it remains a potentially important mechanism.

4.5. Dynamic topography related to mantle insulation and supercontinent cycles Heat energy is mostly lost from the mantle via convection and consequent volcanism at oceanic spreading centers. Formation of supercontinents that persisted for tens of millions of years, covering a wide region of mantle, prevent this method of heat loss, trapping mantle heat produced by radioactive decay in the core. Such mantle insulation may produce a rise in temperature of B20 K throughout the mantle (Gurnis and Torsvik, 1994). Thermal expansion due to this temperature increase produces stress on the base of the lithosphere, creating dynamic topography (Anderson, 1982) with an amplitude of B150 m (Figure 6). As a consequence of this mechanism, continents should experience uplift during supercontinent formation and persistence, followed by subsidence as the supercontinent breaks up and the fragments drift off hot mantle onto adjacent, relatively cool mantle (Anderson, 1982; Gurnis, 1988). Similar effects can be produced by large descending plumes interacting with internal mantle viscosity boundaries (Pysklywec and Mitrovica, 1998). This signature of supercontinent-related dynamic topography is consistent with patterns of Phanerozoic North American cratonic sequence development. The two longest duration lacunae in North America are the base-Sauk and the base-Zuni unconformities (Figure 3). Both formed during periods when North America was part of a supercontinent, Rodinia in the Late Precambrian, and Pangea in the Late Paleozoic and Early Mesozoic. Cratonic erosion and non-deposition would have been accentuated by dynamic topographic highs creating an emergent craton (Sloss and Speed, 1974). Conversely, the three Paleozoic unconformities are of shorter duration (Figure 3), and formed during a period when North America was one of several dispersed continents, overlying cool mantle, and thus having relatively low or submerged (Sloss and Speed, 1974) elevation. Subsidence analysis identiﬁes an anomalously large subsidence event in Late Devonian to Mississippian time, explained by Kominz and Bond (1991) as due to the ﬁnal stages of assembly of Pangea over a dynamic topographic low (Figures 7 and 8). A further striking feature of the base-Zuni unconformity is its east–west asymmetry, reﬂected by the general absence of Mesozoic cratonic strata in eastern North America compared with extensive Jurassic and Cretaceous deposition in the west. This may be in part due to slab-related dynamic topography (see Section 4.4), but numerical modelling (Burgess et al., 1997) suggests that it may also be in part due to continent-scale tilting up-to-the-east. Tilting would have resulted from increasing amounts of uplift eastwards across North America towards the hottest mantle situated beneath the center of Pangea (Figure 6). The extended duration of the sub-Tejas unconformity in eastern North America, and the current elevated topography of North America and the resultant predominantly erosional regime suggest that the dynamic topographic high is somehow persistent, or that elevation is being maintained by some other mechanism, at least locally (e.g. erosional processes on the Blue Ridge escarpment, Spotila et al., 2004). Note that Burke et al. (2003) also alludes to this general mechanism of thermal insulation beneath Pangea (see Figure 2 of Burke et al., 2003) to explain the elevation of Africa and the generation of an apparently atypical number of deep-seated plume events (but see also Coltice et al. (2007) who suggest that thermal insulation alone can generate melting, without any requirement for plumes).

4.6. Mantle downwelling: ‘‘cold spots’’ Hartley and Allen (1994) observed that African intracratonic basins have a distinctive gravity signature, requiring either a cold, dense region in the underlying upper mantle, or a downward-acting dynamic force on the base of the lithosphere. From this they suggested that basins such as the Congo Basin may be situated above convective mantle downwellings, which would generate both the downward force and the elevated density. Such ‘‘cold-spot basins’’ would be expected to have sedimentary ﬁlls developed over a prolonged period of geological time, to lack

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Figure 7 Reduced subsidence curve (Kominz and Bond, 1991) from a stratigraphic section in north-central Iowa. Minimum and maximum curves derived from appropriate compaction corrections. The curves represent components of eustasy, changing continental freeboard and tectonic subsidence. Note the large subsidence event during the Devonian. From Kominz and Bond (1991).

well developed rift precursors, to have low to moderate heat-ﬂow histories and to have an approximately circular planform. Based on this descriptions from Hartley and Allen (1994) a downwelling ‘‘cold-spot’’ mechanism seems to ﬁt several of the observed properties of the North American cratonic basins, perhaps making it a plausible contender to explain at least some aspects of the formation of the basins, as well as some of the uplift and subsidence history on the intervening platforms that presumably overlie relatively hot mantle. Downwelling of this type, persistent for tens of millions of years during the Paleozoic, might explain the long-term subsidence history of cratonic basins more consistently than other mechanisms like thermal re-equilibration after rifting events. However, it might require a relatively stationary position for North America relative to underlying shallow mantle (Burke et al., 2003). More work is required to develop and test this idea.

4.7. Magmatic controls Magmatic underplating is the process of adding igneous melt material, less dense than the asthenosphere, to the base of the lithosphere, causing isostatic adjustment and uplift as a consequence (Brodie and White, 1995). For example, addition of a thickness of 5 km of basalt or gabbro would cause an initial uplift of 625 m, which combined with isostatic response to denudation would lead to a total exhumation of B2.5 km, assuming typical densities for asthenosphere, crust, underplated material and sedimentary rock (Brodie and White, 1995). Although it is a well-understood process, clearly it requires igneous activity to generate signiﬁcant volumes of

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Figure 8 Reduced subsidence curves (Kominz and Bond, 1991) from locations in North America, numbered by location on the map. E¡ects of compaction and sediment loading have been removed. The curves represent components of eustasy, changing continental freeboard and tectonic subsidence. The latter dominates the signal in terms of amplitude. Note the large subsidence event in the Devonian, developed to varying degrees in all 10 curves. From Kominz and Bond (1991).

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melt, some of which seems likely to be extruded at the surface, particularly if there are crustal-scale faults to serve as conduits for rising magma. Therefore, if there is no evidence of surface volcanics associated with uplift events, it is perhaps unlikely that underplating was the cause of the uplift. Another possible magmatic mechanism is intrusion of anorogenic granites that post-date orogenic intrusive events by several million years. Klein (1995) has postulated that intracratonic basin formation was controlled by intrusion of such granites. However, much of the proposed mechanism as described by Klein (1995) appears to have little to do with granite intrusion and more to do with extensional and compressional intraplate stress variations causing rift subsidence and block uplift respectively. Klein (1995) proposes that the granites acted to focus intraplate stress, generating intracratonic basins, but it seems more likely that intraplate stress variations simply reactivated pre-existing tectonic structures created during earlier phases of rifting, orogenesis or general terrane accretion.

5. Phanerozoic Evolution of the Cratonic Platform Cover According to Sloss (1963) the Phanerozoic strata covering the North American craton can be subdivided into a number of stratigraphic sequences. Sequences are deﬁned as ‘‘rock-stratigraphic units of higher rank than group, megagroup or supergroup, traceable over major areas of continent and bounded by unconformities of interregional scope.’’ Students of sequence stratigraphy will recognize this deﬁnition, because it triggered development of many of the later sequence stratigraphic models. Sloss synthesized a large amount of outcrop and subsurface data from the North American continental interior and, through this synthesis, identiﬁed a number of interregional unconformity surfaces that could be traced and correlated across the continent. Interestingly, the interregional unconformities are commonly not obviously distinguishable in their local development from other less areally extensive unconformity surfaces (Sloss, 1963). Based on identiﬁcation of interregional unconformities, Sloss (1963) deﬁned six craton-wide sequences in the Phanerozoic strata of North America, namely the Sauk, Tippecanoe, Kaskaskia, Absaroka, Zuni and Tejas sequences (Figure 3). In general terms, these sequences thin towards and onlap on to cratonic platform areas, and thicken into intracratonic and marginal basins (Figures 4 and 5). The interregional unconformities are of greatest duration in the cratonic interior of the continent, and pass laterally into conformable successions in the marginal basins. Within this broad pattern there is considerable variability in unconformity development. For example, the sequence-bounding unconformities may be less well developed in relatively rapidly subsiding intracratonic basin centers than they are on adjacent uplifted arches (Figures 4 and 5). However, in all cases the strata beneath the unconformity are older than, and the strata overlying the unconformity younger than, the point of maximum ofﬂap in adjacent marginal basins. Burgess et al. (1997) linked the large-scale pattern of cratonic unconformity development to the effects of dynamic topography related to supercontinent formation and break-up (see Section 4.5). Using an improving subsurface data set, the original six Sloss sequences have been further subdivided into subsequences (e.g. Sloss, 1988b). Each subsequence is also bounded by an interregional unconformity, but these may not be present across the entire craton. Subsequences are approximately equivalent to the megacycles of Haq et al. (1987).

5.1. The Sauk sequence (Late Precambrian to Early Ordovician) Sauk strata overlie an interregional, in many places angular, unconformity on late Precambrian sediments and older metamorphic rocks. Abundant stratigraphic evidence suggests that this basal-Sauk unconformity represents a buried land surface showing a few tens of meters of relief (Sloss, 1988b). Basal-Sauk strata are remarkable for their uniformity, being composed of compositionally mature pure quartz sandstones, with only a few local exceptions. Flooding of the basal-Sauk land surface probably occurred due to cratonic subsidence, linked to rifting, continental drift and development of dynamic topography during break-up of a late Neoproterozoic supercontinent (Sloss, 1988b; Bally, 1989; Hoffman, 1989; Burgess et al., 1997). Distribution of Sauk strata is shown in Figure 9. Progressive cratonic transgression was interrupted by periods of minor regression, so that the Sauk sequence can be subdivided into three subsequences (Sauk I, II and III) that progressively overstep and onlap onto the cratonic platform. On the cratonic margins, Late Precambrian and Early Cambrian (Sauk I) strata are basal sandstones passing upwards to interbedded sandstones and shales, capped by thick carbonate strata. Younger Middle Cambrian (Sauk II) strata on the platform are predominantly dolomitized carbonates, with sandstone tongues common near the margin of the Canadian shield. Middle to Late Cambrian (Sauk III) strata, dominated by carbonates, overstep

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Figure 9 Sauk sequence (latest Proterozoic to Early Ordovician) net subsidence rates. Especially signi¢cant labelled features are 1, Cordilleran shelf basin; 3, Appalachian shelf; 4, Ouachita margin; 5, Marathon margin; 8, Mississippi River-Reelfoot rift system; 9,Transcontinental arch;10, Michigan basin;13,Williston basin. From Sloss (1988b). Note that areas of accumulation with rates less than 5 m Myr�1 are not distinguished from locations with no preserved Sauk strata.

earlier Sauk deposits and onlap Precambrian basement on the Transcontinental Arch and the cratonic shield. Topographic features generated by resistant Precambrian lithologies were emergent during much of Cambrian time, but were subsiding and being gradually buried. Final burial of most of the Transcontinental Arch (Figures 1 and 5) did not occur, however, until latest Mississippian time. The North American intracratonic basins ﬁrst developed during deposition of the Sauk sequence, although in some cases Proterozoic precursor rift basins underlie them. Initiation of the Illinois and Michigan basins occurred during Middle Cambrian (Sauk II) time, and the Williston basin became established during Middle-Late Cambrian (Sauk III) deposition. Sauk sequence deposition ended because of relative sea-level fall towards the end of Early Ordovician time. This relative sea-level fall was probably triggered by a change in stress ﬁelds due to plate convergence, and creation of the Laurentian active margin during the Taconic orogeny, but possibly also had a component of glacio-eustastic fall related to ice sheet buildup (Leighton and Kolata, 1990).

5.2. The Tippecanoe sequence (Middle Ordovician to Early Devonian) Distribution of Tippecanoe strata is shown in Figure 10. Erosion of Sauk sequence strata beneath the subTippecanoe unconformity is variable across the craton. In some places, commonly on arches, Tippecanoe strata overstep Sauk strata and lie directly on Precambrian basement (Figures 4 and 5). Locally, dissolution weathering of Sauk carbonates created karst topography, later buried by Tippecanoe strata. The Tippecanoe sequence is divided into two subsequences by a disconformity, most likely related to the long-term glacio-eustatic fall caused by the latest-Ordovician earliest-Silurian Gondwanan glaciation (Sloss, 1988b). There is no evidence of angular discordance on the disconformity, apart from in the Hudson Bay basin, and the tectonic setting appears similar in both subsequences.

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Figure 10 Tippecanoe sequence (Mid Ordovician to Early Devonian) net subsidence rates. Signi¢cant labelled features are 3, Mississippi Valley graben; 4, Illinois Basin; 5, Cordilleran shelf basin; 6, Appalachian shelf; 7, Cincinnati arch;10, Michigan basin; 11,Wisconsin arch;12, Kankakee arch; 13,Williston basin; 15,Transcontinental arch.

Basal Tippecanoe strata exhibit onlap onto the craton, ranging in age from Middle Ordovician (Tippecanoe I) strata in marginal basins, to Silurian (Tippecanoe II) strata in places on the margin of the Canadian Shield and the Transcontinental Arch. This relationship indicates that the Tippecanoe transgression took tens of millions of years to complete and that the shield and cratonic arches were areas of positive relief relative to much of the rest of the cratonic platform, as they had been during Sauk deposition. The Tippcanoe Sequence is absent over much of the western Canadian craton margin (Miall, Chapter 5). Even in the marginal basins, the subTippecanoe unconformity surface has a relief of up to 50 m. The sequence developed under the inﬂuence of compressional collisional tectonics on the eastern cratonic margin due to subduction and collision during the Taconic orogeny. The eastern passive margin became a foreland basin at this time, formed by ﬂexural loading (see Lavoie, Chapter 3 and Ettensohn, Chapter 4). Tectonic differentiation of the craton was initially low, similar to conditions prevailing during Sauk deposition, but increased during Tippecanoe deposition, so that various arches and basins became increasingly well deﬁned, evolving towards a pattern that was then typical of Mid-Paleozoic time (Sloss, 1988b). Siliciclastic rocks are present throughout the U.S. cratonic area at the base of the Tippecanoe sequence, forming a sheet of pure compositionally mature quartz arenite, probably sourced from weathering of crystalline shield areas. Volumetrically, however, the sequence is dominated on the craton by carbonate strata that include numerous reef systems. Thick Silurian evaporite strata are also common, particularly in the intracratonic basins (see Sections 6.1–6.4). Tippecanoe strata covered many arches, much of the western craton, and much of the Canadian shield. Evidence for this includes patterns of onlap and overstep, lack of marginal facies at the limit of present Tippecanoe outcrop, various outliers, and presence of Ordovician and Silurian carbonates preserved as diatreme xenoliths within igneous intrusions in shield areas (Sloss, 1988b; Cecile and Norford, 1993). The extent of Tippecanoe strata is signiﬁcant because it suggests that the Tippecanoe marine transgression culminated in marine ﬂooding of much of the craton, both platform and shield areas, indicating a period of exceptionally elevated relative sea-level. This may have been due, in part at least, to dynamic topography; because at this time the continents were dispersing, moving from dynamic topographic highs to dynamic topographic lows, and therefore undergoing relative sea-level rise (Burgess et al., 1997) (see Section 4.5). This is also consistent with the

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relatively short duration unconformities separating the Sauk, Tippecanoe, Kaskaskia and Absaroka sequences (Figure 3).

5.3. Kaskaskia sequence (mid-Early Devonian to Late Mississippian) The basal Kaskaskia unconformity truncates strata of all older ages, from Sauk through Tippecanoe, demonstrating that there had been signiﬁcant amounts of differential uplift and erosion on the unconformity surface (e.g. Bunker et al., 1988) (Figures 4 and 5). Various arches were reactivated, presumably under a compressional Acadian orogeny stress regime (Figure 3), related to ﬁnal closure of Iapetus, and the landscape was dissected by ﬂuvial channels (Leighton and Kolata, 1990). As with older sequences, the basal Kaskaskia transgression took several million years to re-ﬂood the entire craton, although initial subsidence, ﬂooding and relative sea-level rise were rapid (Figures 7 and 8) (Kominz and Bond, 1991). The latter authors interpreted the rapid subsidence event, and the overall dominance of compressive stress, to indicate the initial assembly of continents over a dynamic topographic low produced by cold mantle (see Section 4.5), an interpretation supported by Burgess et al. (1997). Distribution of Kaskaskia strata is shown in Figure 11. Kaskaskia sequence strata are subdivided into two subsequences, Kaskaskia I ranging from mid-Early Devonian to latest Devonian, and Kaskaskia II, from latest Devonian to Late Mississippian. Kaskaskia I strata progressively onlap basement structures, but in some cases are absent from basement highs (e.g. the Transcontinental Arch, Figures 4 and 5) that were subaerially exposed throughout the Devonian period and only reﬂooded in Mississippian time when they were draped with Kaskaskia II strata. Basal Kaskaskia I sandstones occur locally and are either reworked Sauk and Tippecanoe material or derived from the exposed Canadian Shield. Much of the rest of the subsequence on the craton is dominated by thick bioclastic and biohermal carbonates that pass laterally into black shales and cherts in the marginal basins. Evaporite units within the carbonates represent periods of sabkha formation during periodic regressions, and carbonate deposition was also interrupted occasionally by periods of anoxia, generating black shales. The unconformity separating Kaskaskia I from Kaskaskia II strata was marked by uplift and erosion on arches and domes, but basinal areas suffered only a minor hiatus. However, Kaskaskia II strata record onset of signiﬁcant siliciclastic inﬂux from the Appalachian, Ouachita, Marathon and Antler orogenic belts forming on the continental margins in the east, south and west (Figure 3), and from the partly re-emergent Canadian Shield. Initial Kaskaskia II deposition shows evidence of restricted anoxic conditions but this changed as transgression progressed, circulation improved and carbonate deposition was established (Sloss, 1988b). Previously emergent basement features were transgressed and buried. Orogenic activity increased during the Late Mississippian, previously dominant carbonate production across the craton was ﬁnally smothered, and subsequent deposition was dominated by siliciclastic material derived from orogenic margins. Consequently in many areas the youngest Kaskaskia strata preserved are sandstones and shales, in places forming cyclothems capped by coal.

5.4. Absaroka sequence (Late Mississippian to Early Jurassic) The sub-Absaroka unconformity records extensive denudation that removed up to thousands of meters of older strata, particularly from old and new areas of positive relief (e.g. the Transcontinental Arch). In places, denudation removed all older Phanerozoic strata, so that basal Absaroka units rest directly on Precambrian basement. Much of this cratonic tectonic activity can be related to orogenic events on the western, southern and eastern cratonic margins creating the Antler, Ouachita-Marathon and Alleghenian orogens, respectively (Sloss, 1988b; Leighton and Kolata, 1990; see also Miall, Chapter 8; Ingersoll, Chapter 11). Compressive stress from orogenic events transmitted into the cratonic interior reactivated numerous basement structures causing uplift and erosion. Local uplifts on the margins of the intracratonic basins left them isolated from the rest of the platform. Leighton and Kolata (1990) postulated that the widespread uplift due to compressive stress ﬁelds may have been related to collision of Gondwana and Laurussia along the Eurasian-African margin. Given the icehouse setting, the degree to which a glacio-eustatic component also inﬂuenced development of the unconformity is debatable (Beuthin, 1994; Ettensohn, 1994). Distribution of Absaroka strata is shown in Figure 12. As in older sequences, the basal Absaroka transgression was prolonged, generating upper Mississippian to Triassic transgressive units (Sloss, 1963). Once again, the oldest Absaroka strata are found in the marginal basins, and the youngest strata drape cratonic arches. In contrast to the older sequences, Absaroka cratonic strata are dominantly siliciclastic. Much of this material was derived from within the craton, either from the shield or from emergent arch areas, but marginal orogenic topography contributed sediment to the edges of the craton. The Absaroka sequence is split into three subsequences, spanning Latest Mississippian to Early Permian (Absaroka I), Middle to Upper Permian (Absaroka II) and Triassic to Early Jurassic (Absaroka III) (Sloss, 1988b).

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Figure 11 (A) Kaskaskia I subsequence (mid-Early Devonian to Late Devonian) net subsidence rates. Signi¢cant labelled features are 2, Cincinnati arch; 6, Illinois basin; 10, Michigan basin; 12, Catskill siliciclastic wedge. (B) Kaskaskia II subsequence (latest Devonian to Late Mississippian) net subsidence rate. Signi¢cant labelled features are 1,Williston basin; 3, Reelfoot-Illinois basin; 4, Cordilleran foreland basin; 6, Michigan basin; 8, Black Warrior basin; 9, Appalachian foreland basin.

Figure 12 (A) Absaroka I subsequence (latest Mississippian to Early Permian) net subsidence rates. Signi¢cant labelled features are 1, Ouachita margin; 2, Marathon margin; 3, Uncompahgre uplift; 10, Paradox basin; 11, Eagle basin; 12, Denver basin; 13, Anadarko basin; 14, Fort Worth basin; 16, Delaware basin; 18, Nemaha uplift; 21, Michigan basin; 22, La Salle anticline; 23, Cincinnati arch; 26, Central Kansas uplift; 22, Illinois basin; 30, Appalachian basin. (B) Absaroka II subsequence (Middle and Late Permian) net subsidence rates. Signi¢cant labelled features are 1, Delaware basin; 2, Midland basin; 3, Central Platform basin; 4, Anadarko basin; 5, Alliance basin; 6,Williston basin; 7, Colorado Plateau area. (C) Absaroka III subsequence (Triassic to Early Jurassic) net subsidence rates. Signi¢cant labelled features are 1, Extensional rift basins; 2, Permian basin; 3,Wiliston basin; 4,Western limit aeolian sandstones.

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The subsequences are generally deﬁned on the basis of relatively minor disconformities; generally, the Absaroka sequence records a steady shallowing-upward trend. However, Bally (1989) noted that Absaroka I strata were inﬂuenced by a stress regime related to the Ouachita orogeny (Figure 3), while Absaroka II deposition was dominated by the Alleghenian orogeny (Figure 3; see Miall, Chapter 8). Both these orogenies occurred during assembly of the Pangea supercontinent. In contrast, Absaroka III deposition marked the onset of the subsequent break-up of Pangea. Absaroka I strata in southern, western and midcontinent areas drape a complex mosaic of uplifted and eroded fault blocks and arches (Figure 12; see Blakey, Chapter 7), with proximal coarse siliciclastic material and evaporitic strata generally forming more distally. Delta systems prograded southwestward on the eastern part of the craton, sourced from the Appalachian and eastern Canadian Shield areas. Maximum cratonic transgression occurred in Pennsylvanian time establishing predominantly marine conditions in the southwest, and passing into more terrestrially dominated deltaic strata in the north-east, perhaps inﬂuenced by glacio-eustatic and local tectonic factors to create cyclothems (Miall, Chapter 8). Absaroka II strata indicate considerably less tectonic differentiation (Figure 12), with deposition of alternating marine and continental strata across a broad, apparently stable cratonic area (Sloss, 1988b). Siliciclastic sediment inﬂux from the northeast continued, but the southern and western cratonic areas were dominated by carbonate deposition. The trend of southwestward retreat of the sea was continued during deposition of Absaroka III strata. This subsequence records a ﬁnal episode of cratonic deposition in mostly terrestrial conditions and limited to the western margin of the craton (Figure 12). Absaroka deposition was ﬁnally terminated by relative sea-level fall and development of a craton-wide erosional unconformity surface, cutting into Pennsylvanian strata in the east, and Permian to Lower Jurassic strata in the west. The prolonged regression during later stages of Absaroka deposition, and the consequent long midcraton hiatus before onset of Zuni deposition are consistent with gradual elevation and up-to-the-east tilting of North America over a dynamic topographic high generated by supercontinent insulation (Burgess et al., 1997).

5.5. Zuni sequence (Middle Jurassic to Early Paleocene) Deposition of Zuni strata, ranging in age from Middle Jurassic to Early Paleocene, was initially limited to the Cordilleran margin in western North America (Figure 13; see Miall et al., Chapter 9). The sequence then records progressive overstep on to the craton into Late Jurassic and Cretaceous time, and onset of deposition in the Gulf and Atlantic rift-margin basins. Overstep peaked in the Late Cretaceous, with deposition of marine and terrestrial strata over much of the cratonic platform and the Canadian Shield, indicated by signiﬁcant preserved thickness, or by remnant outliers (Bunker et al., 1988; Sloss, 1988b). This pattern is similar to that developed during Sauk deposition in Cambro-Ordovician time (Sloss, 1988b) and probably represents somewhat similar tectonic conditions, with deposition occurring during cratonic subsidence driven by evolving dynamic topography related to ongoing supercontinent break-up (Burgess et al., 1997). The vast extent of Cretaceous deposition is an indication of the extent of the Cretaceous Interior Seaway that ﬂooded much of North America (see Miall et al., Chapter 9 on Western Interior basins). This is, in part at least, due to the low-angle penetration of the Farallon slab beneath western North America and the resulting dynamic topographic low, best developed on the Cordilleran margin, but extending across much of the craton (Burgess et al., 1997; and see Section 4.4). High global sea-level no doubt also played a part in generation of the Cretaceous seaway, but thickness trends and longwavelength tilting of the strata indicate the inﬂuence of slab-related dynamic topography (Burgess et al., 1997; Burgess and Moresi, 1999). Division of the Zuni into three subsequences (Sloss, 1988b) indicates that the pattern of gradual overlap during transgression of the Interior Seaway was interrupted twice, once in the Early Cretaceous Epoch, separating Zuni I and Zuni II subsequences, and again at the beginning of the Late Cretaceous, separating Zuni II and Zuni III subsequences. In both case relative sea-level falls caused terrestrial deposition, subaerial erosion and hiatus. The Zuni sequence is thus subdivided into three subsequences. The hiatus at the base of the Zuni II subsequence was particularly signiﬁcant, representing around 10 Myr of non-deposition, and having an equivalent unconformity developed throughout Europe (Leighton and Kolata, 1990). Consequently this subsequence boundary is interpreted as formed by a prolonged eustatic low. During Early Zuni I deposition the cratonic interior was the main source of siliciclastic detritus, but orogenic loads developed in the west became the dominant sediment source from Late Jurassic time onwards (Sloss, 1988b). Throughout Zuni deposition, large volumes of siliciclastic sediment were shed from the developing mountains, extending 300 km or more eastwards (Miall et al., Chapter 9). Further east, thick marine shales and carbonates accumulated across much of the cratonic platform. Subsidence rates were eventually outpaced by sediment supply rates, and perhaps by eustatic fall, and ﬁnally terminated by a change to tectonic uplift (Leighton and Kolata, 1990). Consequently Paleocene platform deposition was mostly terrestrial, and Zuni deposition was

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Figure 13 (A) Zuni I subsequence (Middle Jurassic to earliest Cretaceous) net subsidence rate. (B) Zuni II subsequence (midEarly Cretaceous to earliest Late Cretaceous) net subsidence rates. (C) Zuni III subsequence (Late Cretaceous to Early Paleocene) net subsidence rates. Signi¢cant labelled features are 1,Wind River uplift; 2, Front Range uplift; 3, Kaibab uplift; 4, Red Desert-Hanna basin; 5, Piceance-Washakie basin; 6, Denver basin; 7, Pedregosa basin.

terminated by development of a major unconformity, often showing angular discordance, indicative of signiﬁcant tectonic tilting.

5.6. Tejas sequence (Late Paleocene to present) Development of the sub-Tejas unconformity in the west can largely be related to deformation of crustal blocks during the Laramide orogeny (Sloss, 1988b; Lawton, Chapter 12). Elsewhere on the craton a combination of low eustatic sea-level and long-wavelength uplift can be invoked (Burgess et al., 1997; Burgess and Moresi, 1999). Deposition of Tejas strata was areally more limited than in previous sequences. Laramide crustal blocks and associated small basins ﬁlled with Late Paleocene to middle Eocene Tejas I ﬂuvial and lacustrine deposits. Tejas II

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and Tejas III strata consist of large volumes of ﬂuvial strata deposited in the west, shed from and burying the Laramide Mountains. Although cratonic Tejas volumes are negligible compared to strata from previous sequences, non-cratonic Tejas strata form thick successions in the Gulf Coast and Atlantic passive margin basins (McCracken, Chapter 14; Galloway, Chapter 15). The area of western North America previously occupied by Zuni and Tejas depocenters was uplifted during Neogene times to elevations of 1.5 km or more, and the Zuni and Tejas strata tilted, exposed and subjected to erosion. This process can be explained reasonably well by the decay of a dynamic topographic low associated with the Farallon slab. The dynamic topographic low was responsible in part for accommodation creation in the Zuni and Tejas depocenters. During the ﬁnal stages of subduction and slab detachment the dynamic topography was reversed and rock and surface uplift occurred (Mitrovica et al., 1989; Burgess et al., 1997). The current elevated hypsometry of the rest of the North American craton is more difﬁcult to explain, but given similar hypsometries on other continents, may be caused by some long-wavelength mantle effect, perhaps elevated temperatures persisting from Pangean mantle insulation. Relatively low oceanic water volumes due to signiﬁcant water storage in polar ice caps also contributes to the observed elevated hypsometry.

6. The North American Intracratonic Basins The large-scale geometry and stratal architectures of the four intracratonic basins discussed here are summarized in Figure 14. Their increased stratal thickness relative to surrounding platform areas, and their spatial distinction from marginal basins are shown in Figure 2.

6.1. The Michigan Basin The Michigan basin is an oval-shaped intracratonic depocenter covering most of the state of Michigan. It contains around 4,500 m of strata composed mostly of carbonate and evaporite rocks with only subordinate siliciclastic sedimentary rocks (Fisher et al., 1988) (Figure 14). The basin is situated atop a Precambrian failed rift system (Nunn et al., 1984). Phanerozoic strata within the basin range in age from Cambrian to Jurassic, and represent the Sauk to Zuni sequences of Sloss (1963) (Figure 15). 6.1.1. Phanerozoic history of the basin Deposition in the basin began in Late Cambrian time, ﬁlling a Precambrian structural low with up to 2,000 m of sandstone, shale and sandy dolomites (Fisher et al., 1988) forming a series of deepening-upward cycles (Catacosinos et al., 1990). These initial deposits form the Sauk II sequence, and unconformably overlie Precambrian rocks. Sauk II strata pass conformably up into overlying strata of the Sauk III sequence. In this case the Sauk II–III subdivision is based on biostratigraphic correlation of lithostratigraphic units with other areas of Cambrian strata in North America. Sauk III strata are dominated by shallow-water and peritidal dolomites (Catacosinos et al., 1990) with subordinate sandstones and lateral transitions into shale towards the basin center. Deposition of the Sauk III sequence was terminated by a developing unconformity ranging in age across the basin from Late Cambrian to Middle Ordovician. Erratically distributed basal units of the overlying Tippecanoe I sequence suggest deposition over a highly irregular surface, perhaps a karst terrain developed on underlying Sauk III carbonates. The sequence shows a general ﬁning-upwards trend, from basal nearshore marine sandstone units, into overlying shale and shallowmarine platform carbonate strata. During deposition of this sequence, the basin departed from its previous bulls-eye pattern and tilted downwards to the east for a period of 10–15 Myr. This tilting event has been ascribed to subsidence due to dynamic topography forming above a slab of oceanic lithosphere being subducted westwards beneath the craton from the Iapetus margin to the east (Coakley and Gurnis, 1995). Tippecanoe I deposition was terminated at the end of the Ordovician period by development of a basin-wide unconformity. Deposition recommenced in Silurian time, forming the Tippecanoe II sequence. Strata in this sequence are predominantly carbonates and evaporites, and include abundant patch reefs, as well as a barrier reef carbonate, up to 210 m thick, that encircles the basin (Fisher et al., 1988; Catacosinos et al., 1990). Strata younger than Tippecanoe II are absent on the northern edge of the basin, and in the rest of the basin Tippecanoe II strata are overlain unconformably by the Kaskaskia I sequence. Kaskaskia I strata are dominated by shelf carbonates, shale and sabkha evaporites. The upper parts of the sequence include black shales interﬁngering with less organic-rich gray green shales. The top of the sequence is marked by presence of shale interpreted as pro-delta deposits, formed by a river ﬂowing south from Canada. Kaskaskia II strata overlie Kaskaskia I conformably, but occupy a signiﬁcantly smaller area, due largely to a

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Figure 14 Schematic cross-sections of the four intracratonic basins of North America. The sequences of Sloss (1963) are coded by di¡erent stipple patterns. From Bally (1989) and Leighton and Kolata (1990).

combination of non-deposition and post-Mississippian erosion over structural highs. The sequence is dominated by terrigenous input. Deltaic deposition continued in the east, with more proximal deposits forming on a delta top. In other parts of the basin marine shales and sandstones dominate, with only occasional thin limestones. In total, the Kaskaskia I and II subsequences reach a maximum 5,300 m thick. During latest Mississippian, the Michigan basin was an area of low positive relief, by-passed by large rivers transporting sediment into the Illinois basin to the south (see Section 6.2 on the Illinois basin).

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Figure 15 Chronostratigraphic diagram of strata in the Michigan Basin. From Catacosinos et al. (1990).

Pennsylvanian Absaroka I strata are limited to the basin center, lie disconformably atop the Kaskaskia II strata, and are no more than 230 m thick. Interbedded shales, coals and limestones indicate alternating marine and nonmarine deposition typical of Pennsylvanian strata throughout North America and Europe. Strata forming the Jurassic Zuni I sequence are even more geographically restricted and less than 120 m thick. They have been interpreted to represent ﬂuvial paleovalley ﬁlls, either of Kimmeridgian or Bajocian-Bathonian age. Decreasing depositional area in the basin, either due to post-depositional erosion (in the case of the Absaroka strata), or limited deposition (in the case of Zuni strata), can be interpreted in the wider context of supercontinent formation and uplift related to mantle insulation (see Section 4.5).

6.2. The Illinois Basin The Illinois basin is a saucer-shaped, oval depression ﬁlled with up to 7,000 m of Paleozoic strata ranging in age from Upper Cambrian to Permian (Figure 16), with only minor thicknesses of Mesozoic and Cenozoic strata.

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Figure 16 Chronostratigraphic diagram of strata in the Illinois Basin. From Bushbach and Kolata (1990).

Peter M. Burgess

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Approximately equal volumes of siliciclastic and carbonate lithologies occur. Lower Cambrian strata ﬁll a failed rift beneath the younger basin ﬁll (Bushbach and Kolata, 1990). Basement rocks are Precambrian granite and rhyolite dated as between 1,420 and 1,500 Ma. The Sauk, Tippecanoe, Kaskaskia and Absaroka sequences are well developed in the basin, but Zuni and Tejas strata are poorly represented, apart from a Quaternary cover over much of the basin (Figure 16). 6.2.1. Phanerozoic history of the basin The oldest strata in the basin are Lower and Middle Cambrian sediments ﬁlling the New Madrid Rift complex (Braille et al., 1982; Collinson et al., 1988) in the southern part of the basin. The rift probably formed during the break-up of the Rodinia supercontinent (Braille et al., 1982) and rifting had probably ended by Late Cambrian time. Post-rift subsidence was more widespread, so that the rest of the Sauk sequence was deposited in Late Cambrian and Early Ordovician time over an area extending north away from the rift complex (Figure 10). Potsdam Megagroup strata are predominantly sandstones and form the lower part of the Sauk sequence, present over much of the basin, and sourced predominantly from the shield areas to the north. Knox Megagroup carbonates overlie the sandstones in the south, and interﬁnger in the north and east. During Sauk deposition the basin was open to the south, forming a large embayment on the edge of the craton. Separation of the Illinois and Michigan basin areas by arch uplift began during Sauk deposition. The whole Sauk sequence is interpreted as products of a shallow, subtidal marine system recording a gradual transgression and deepening, terminated ﬁnally by development of the sub-Tippecanoe unconformity surface. The unconformity represents a 10–15 Myr hiatus, and in places up to 31 m of relief is preserved on the unconformity surface. Deposition of the Tippecanoe sequence began in the south and spread northwards. The thickest part of the sequence is in southern Illinois and western Kentucky. Tippecanoe strata are dominated by carbonate rocks, with only limited volumes of siliciclastic deposition recorded. The Tippecanoe I and II subsequences are deﬁned by an unconformity separating Ordovician and Silurian strata, marked by slight erosion (Collinson et al., 1988). Several other hiatuses and minor erosion surfaces indicate more complex relative sea-level history than during Sauk deposition, in part due to local and regional uplift, and in part due to glacio-eustasy (Collinson et al., 1988). Tippecanoe I strata represent predominantly peritidal and shallow-subtidal deposition, with deeper subtidal deposition in the west at the end of the Ordovician. Similar conditions prevailed during Tippecanoe II deposition, but with local reef build-ups, and more siliciclastic input. Most siliciclastic strata were derived from the Ozark uplift, the Transcontinental Arch, and from the Canadian Shield, but some of the Upper Ordovician ﬁne-grained siliciclastic strata originated from an Appalachian source. Erosion on the sub-Kaskaskia unconformity mirrors the distribution of Tippecanoe thicknesses. Least erosion occurred in the deepest parts of the basin in southern Illinois and western Kentucky (Figure 10), and in places deposition was uninterrupted. Greatest amounts of erosion, cutting down to Ordovician or pre-Tippecanoe occurred over the Ozark Dome and the Northeast Missouri Arch (Kolata and Olive, 1990). Various tectonic elements within the basin were uplifted and eroded during unconformity formation (Collinson et al., 1988), and in places solution ﬁssures up to 22 m deep, and channels up to 16 m deep where formed (Devera and Hasenmueller, 1990). Kaskaskia strata are again carbonate dominated, but with an increased clastic input relative to earlier sequences. The sequence is divided into subsequences at the Devonian-Mississippian boundary, but the boundary is largely conformable, identiﬁed only from biostratigraphic data (Collinson et al., 1988). Further transgression and deposition of carbonate strata across the basin followed initial deposition of sandstone above the basal unconformity in the basin center. These strata were in turn overlain by laminated black shales (New Albany Group), deposited across large portions of the craton (see Section 5 for an overview of Sloss sequences), and indicating extensive anoxic marine conditions. Carbonate deposition in the basin resumed, forming the Mammoth Cave Megagroup (Figure 16). Mixed siliciclastic and carbonate deposition occurred in the latest Mississippian when delta systems associated with a large river system prograded into the basin, sourced either from the Canadian shield or from the adjacent Appalachian mountains (Collinson et al., 1988). Onset of clastic deposition is probably linked to development of the Alleghenian and Ouachita orogenies. Kaskaskia deposition terminated as the basin became subaerially exposed, karst topographies formed and ﬂuvial systems became deeply incised. Maintaining the pattern of earlier unconformities, the sub-Absaroka unconformity decreases in duration southwards to the deepest parts of the basin. The entire basin area was subaerially exposed, with contemporaneous tectonic activity causing surface uplift and accentuating erosion so that ﬂuvial valleys up to 130 m deep and 32 km wide were formed (Collinson et al., 1988). Early Absaroka deposition was marked by continued tectonic activity and intra-basinal highs either being eroded, or accumulating relatively shallow-water deposits. Deposition commenced in the south where over 700 m of sediment accumulated, and transgression

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deposited strata progressively northwards. Absaroka strata are dominantly siliciclastic, formed in ﬂuvial, deltaic and shallow-marine environments. The basin opened to the south and had a low gradient, so that small ﬂuctuations in relative sea-level caused large lateral shifts in depositional environment, reﬂected in numerous cycles of delta progradation and subsequent marine ﬂooding. Deposition was continuous through into Permian times in parts of the basin, and coal maturity and shale compaction indicates that certain areas of Pennsylvanian strata were covered by signiﬁcant thicknesses of Permian strata, since eroded. There is no evidence of deposition younger than Early Permian until Cretaceous strata of the Zuni sequence were deposited. Thus the base-Zuni unconformity represents a hiatus of approximately 200 Myr, probably associated with whole-continent tilting upto-the-east by dynamic topography related to supercontinent insulation (Burgess et al., 1997; see Section 4.5). Various faults and arches were active during the sub-Zuni hiatus, and igneous rocks of Permian age occur in southern Illinois (Collinson et al., 1988). Active tectonic features were eroded and reduced to base level during the hiatus, so that the ﬁrst Zuni strata in the Late Cretaceous epoch are non-marine, overlap eroded Ordovician to Mississippian strata and in some areas overlie a basal soil (Kolata and Olive, 1990). Zuni strata reach a maximum thickness of around 600 m in the southern part of the basin, and are poorly preserved elsewhere, with only scattered outliers in western Illinois preserving a maximum thickness of 30 meters. These outliers are important, however, because they represent some of the easternmost examples of Upper Cretaceous strata, and were deposited in a nearshore marine environment. They were therefore probably deposited on the eastern margin of the Cretaceous Interior Seaway (Kolata and Olive, 1990), in accommodation generated by the easternmost limits of slab-related dynamic topography (see Section 4.4). Petrological data suggest a source in the metamorphic terranes of the Appalachian area to the east. Unconformities within the Zuni strata suggest numerous oscillations in relative sea-level; depositional environments range from terrestrial to shallow marine. Marine Paleocene strata overlie Cretaceous strata across an unconformity and mark the ﬁnal deposition of the Zuni sequence. Tejas strata in the Illinois basin are Pliocene in age, and consist of terrestrial gravel. The basin area thus remained emergent throughout Cenozoic time after an initial period of Paleocene marine deposition. Latest Tejas deposition consists of Quaternary glacial deposits, with notably low preservation potential.

6.3. The Williston Basin The Williston basin is another elliptical intracratonic depocenter, ﬁlled with up to 4,900 m of strata spanning the whole of Phanerozoic time (Figure 17). Upper Cambrian strata initiated the basin ﬁll and it was initially a cratonic-margin basin, only becoming intracratonic during the Cordilleran orogeny as material was accreted to the western continental margin. Data on basement rocks are rather sparse, but Archean Canadian shield rocks extend beneath the basin, and there is also evidence for island arc and oceanic crust basement to the west of the shield rocks (Green et al., 1985; Gerhard et al., 1990). There is abundant evidence for inﬂuence of faults and arches active throughout Phanerozoic time and inﬂuencing stratal architectures. 6.3.1. Phanerozoic history of the basin Deposition of Upper Cambrian Sauk strata began when the basin was still only an embayment on the western cratonic margin. Thinning of Sauk strata across structural features suggest that there was signiﬁcant relief on the base-Sauk unconformity surface. Sauk strata are dominantly quartzites and siliciclastic conglomerates, passing upwards into conglomerates with carbonate clasts. Sauk deposition terminated as the Taconic orogeny began on the eastern cratonic margin, perhaps suggesting some genetic link caused by intraplate stress variations. Although the ﬁrst B120 m of Tippecanoe strata are marine siliciclastic rocks, formed during initial transgression of the sub-Tippecanoe unconformity surface, the rest of the sequence (up to B650 m thick) is dominated by carbonate and evaporite strata. The carbonates are mostly shallow-marine and peritidal limestones and dolomites, interﬁngered with evaporite strata formed during intervals of reduced basin circulation. Various structural elements within the basin became established during or at the end of Tippecanoe deposition, some with signiﬁcant relief inﬂuencing depositional patterns. The Tippecanoe sequence is terminated by a basin-wide unconformity surface showing signiﬁcant karst development on underlying carbonate strata. As in other North American basins Kaskaskia strata can be subdivided into Devonian (Kaskaskia I), and Mississippian (Kaskaskia II) successions, based on a regional, though not basin-wide unconformity (Figure 17). The unconformity is particularly well developed atop structural features, indicating differential uplift, and is absent in the basin center (Gerhard et al., 1990). Both subsequences are carbonate and evaporite dominated, with only minor, ﬁne-grained siliciclastic strata present, in contrast to the pattern seen in other intracratonic basins in the east. The Williston basin underwent renewed subsidence and the marine connection to the Alberta basin in the west was cut. Both events were probably related to development of subduction and a foreland basin system, the Elk Point basin, to the west during the Ellesmerian orogeny (Sloss, 1988b; Leighton and Kolata, 1990).

Phanerozoic Evolution of the Sedimentary Cover

Figure 17

57

Chronostratigraphic diagram of strata in the Williston Basin. From Gerhard et al. (1990).

Kaskaskia II strata show abundant evidence of ﬂuctuating relative sea-level, with rapid transgression followed by gradual progradation (Gerhard et al., 1990). Basin-margin strata were deposited in shallow-marine and peritidal settings, and subjected to extensive subaerial exposure prior to ﬁnal burial. Basin center strata comprise a

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series of shallowing-upward carbonate cycles, capped by crinoidal bioclastic banks and sabkha facies. Following the pattern of increasing structural inﬂuence, the post-Kaskaskia unconformity is present across the basin but best developed on various uplifted arches. Tectonic inﬂuence on the unconformity development is indicated by contemporaneous uplift in nearby areas such as the Canadian Shield (Aitken, 1993) and the Souix Arch in South Dakota. This also manifests itself in a change of depositional style in the Absaroka sequence when siliciclastic strata derived from these uplifted extrabasinal sources became dominant, with the main sediment supply from the south. Depositional environments consisted of terrestrial siliciclastic systems interﬁngering with basinal marginal-marine siliciclastic and evaporite environments. These marginal-marine systems prograded during deposition of the Absaroka strata, and the basin became increasingly restricted and saline, with more basin-center evaporites being deposited. The ﬁnal phase of Absaroka deposition was marked by a decrease in salinity and deposition of carbonates followed by a ﬁnal basin ﬁll of ﬁne to medium grained siliciclastic strata. Deposition was ﬁnally terminated by erosion, presumably triggered by a new phase of relative sea-level fall. By this time, accretion on the active margin to the west had left the basin in a truly intracratonic setting. Zuni strata are Jurassic to Tertiary in age, and consist of mixed carbonate and siliciclastic rocks. Deposition was not continuous; Jurassic and Cretaceous strata are separated by an unconformity representing a 26 Myr hiatus (Gerhard and Anderson, 1988). Jurassic deposition was more carbonate dominated, but siliciclastic strata became more prevalent into the Cretaceous when the area was part of the Cretaceous Interior Seaway, accumulating a series of transgressive and regressive cycles characteristic of the succession at this time (Miall et al., Chapter 9). The rest of the Zuni sequence, and the subsequent Tejas sequence follow the patterns observed in the Interior Seaway, summarized in Chapter 9.

6.4. The Hudson Bay Basin The Hudson Platform is comprised of a series of basins with intervening arches (Figure 18). The largest of these basins is the Hudson Bay basin. The basin was initially a relatively simple saucer-shaped depression, but evolved a more complex geometry later in the Paleozoic when pre-existing fault systems and arches were reactivated by development of regional stress ﬁelds. Basement rocks consist of ensialic Archean protocontinent in the west and Lower Proterozoic fold-belts, including ophiolites, in the east (Roksandic, 1987). The preserved basin ﬁll is mostly limited to Ordovician to Devonian strata, though there are minor areas of Cambrian and Cretaceous strata present (Figure 19). Total composite thickness does not exceed 2,500 m (Figures 14 and 18). 6.4.1. Phanerozoic history of the basin During the Cambrian period the Canadian Shield area was still topographically high and therefore remained emergent. The transgressive episode recorded in Sauk strata across much of the U.S. portion of the North American craton had little recorded effect in the Hudson Bay area, with Sauk strata present only in small areas in the south-east. These Sauk strata consist of c.60 m of unfossiliferous orthoquartzite sandstones and conglomerates, overlain by sandy and stromatolitic dolostones (Sanford, 1987). From Middle Ordovician time, marine transgression was more pronounced, ﬂooding the basin and onlapping strata onto the Transcontinental Arch (Figure 18). Approximately 180 m of strata accumulated, directly overlying Precambrian basement. A middle Ordovician regression ended with renewed transgression, and deposition of between 50 and 200 m of mixed carbonate and siliciclastic strata, once again onlapping and thinning onto the Transcontinental Arch to the west. Limited evidence suggests that the transgression may have peaked with a marine connection through to the St. Lawrence Platform (Figure 1). The trend of increasing transgression and encroachment of the Hudson Platform, Canadian Shield and the Transcontinental Arch continued during deposition of Ordovician to Silurian Tippecanoe strata, representing the continued inﬂuence of a developing dynamic topographic low formed as the continents dispersed (see Section 4.5). Marine connections were established across the whole platform and as far as the Arctic and St. Lawrence platforms. During this interval the Hudson Bay basin ﬁrst became an identiﬁable depocenter, distinct from the surrounding platform area, suggesting operation of a relatively focussed subsidence mechanism in the area of the basin. A hiatus separates Ordovician from Silurian strata, deﬁning the Tippecanoe I and Tippecanoe II subsequences. Ordovician and Silurian lithologies include black bituminous oil shales, and carbonates that formed thick basin-margin successions, including extensive barrier reef units that ringed the basin and fringed adjacent tectonic arches (Sanford, 1987). Conditions for development of these Middle Silurian reef systems may have been aided in part by reactivation of the various structural arches underlying the central portion of the basin. Reactivation caused uplift, initially producing shoaling in the central basin area, and ﬁnally generating an emergent arch, facilitating reef growth, and dividing the basin into eastern and western sub-basins

Figure 18 East-west restored sections: Canadian craton. CO1, Early Cambrian to Canadian; O1,Whiterockian to Blackriverian; O2,Trentonian to Edenian; OS1, Maysvillian toWenlock; SD1, Ludlow to Gedinnian; D1, Seigenian to Eifelian; D2, Givetian to Famennian. From Sanford (1987).

60

Figure 19

Peter M. Burgess

Chronostratigraphic diagram of the Hudson Platform area. From Roksandic (1987).

(Roksandic, 1987). Uplift of surrounding cratonic areas left the Hudson Platform as an inland sea with only narrow connections to the Arctic Platform and to the Hudson Straits in the north-east. Kaskaskia strata are separated from underlying Tippecanoe strata by an Early Devonian unconformity surface (Roksandic, 1987). However, Kaskaskia strata continue many of the trends and characteristics of Tippecanoe strata, although Kaskaskia strata are not present in the northeast of the basin. The arch running through the center of the basin was progressively buried during Middle and Late Devonian time (Figure 18). Some initial Kaskaskia

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61

deposits developed locally are terrigenous red-bed units that interﬁnger with marine carbonates. Most of the sequence is composed of mixed siliciclastic and carbonate strata, with Upper Devonian black bituminous shales developed in the south-east of the basin, possibly sourced from the nearby Appalachian orogen (Sanford, 1987). Kaskaskia deposition was terminated by marine regression, and the Hudson Platform was subjected to extensive erosion throughout the rest of Phanerozoic time (Figure 19), although it is not clear how much Absaroka strata was deposited during Mississippian and Pennsylvanian time and subsequently eroded. Continental deposits are only preserved locally on the platform (Sanford, 1987). If the uplift in the basin occurred in Mesozoic times, it may have been due to uplift of eastern North America over a dynamic topographic high developed by Pangean mantle insulation (Burgess et al., 1997).

7. Summary 1. The North American craton is not simply an unchanging, stable platform accumulating strata and inﬂuenced only by changes in global sea-level. Rather, viewed on a timescale of tens to hundreds of millions of years, it is a dynamic tectonic environment inﬂuenced by various plate tectonic and mantle processes. 2. The Sloss cratonic sequences record the history of this dynamic tectonic environment. 3. Variations in dynamic topography generated by subducting lithospheric slabs and by thermal insulation of mantle beneath supercontinents can explain much of the large-scale architecture of the Sloss sequences but more detailed plate tectonic reconstructions and associated mantle convection models are necessary to further test and develop these explanations. 4. Intraplate stress clearly played a large role in generating the cratonic sequences by reactivating pre-existing structures and driving subsidence and uplift. Variations in intraplate stress through time can be related to some degree to orogenic and other plate tectonic events occurring on the cratonic margins. 5. Given present available evidence and theory, the North American cratonic basins seem most likely to be due to a combination of mantle downwelling and focused intraplate stress variations, in some cases with an element of long-wavelength tilting due to subduction-induced dynamic topography, and in some cases with an initial trigger by lithospheric stretching. This is by no means a deﬁnitive conclusion, however. Much work remains to be done to test, and to conﬁrm or refute these ideas.

ACKNOWLEDGMENTS Thanks to Mike Gurnis who nurtured my interest in the North American cratonic sequences and the inﬂuence of mantle dynamics, and a posthumous acknowledgement to Larry Sloss for his monumental and groundbreaking work in describing, deﬁning and understanding North American cratonic sequences. Thanks to Andrew Barnett, Wes Gibbons, Octavian Catuneanu and Andrew Miall for reading and improving previous drafts of this work.

REFERENCES Aitken, J. D., 1993, Tectonic evolution and basin history; Chapter 5, in Stott, D. F. and Aitken, J. D. eds., Sedimentary cover of the craton in Canada, Geological Society of America, The Geology of North America, v. D-1, pp. 483–502. Allen, P. A., Verlander, J. A., Burgess, P. M., and Audet, D. M., 2000, Jurassic giant erg deposits, ﬂexure of the US continental interior, and the timing of the onset of cordilleran shortening. Geology, v. 28, pp. 159–162. Anderson, D. L., 1982, Hotspots, polar wander, Mesozoic convection and the geoid. Nature, v. 297, pp. 391–393. Bally, A. W., 1980, Basins and subsidence — a summary, in Bally, A. W., Bender, P. L., McGetchin, T. R. and Walcott, R. I. eds., Dynamics of plate interiors, American Geophysical Union, Washington, DC, USA, v. 1, Geodynamical Series. Bally, A. W., 1989, Phanerozoic basins of North America, in Bally, A. W. and Palmer, A. R. eds., The geology of North America-an overview, Geological Society of America, Boulder, CO, v. A, pp. 397–446. Beaumont, C., Quinlan, G. M., and Hamilton, J., 1987, The Alleghanian Orogeny and its relationship to the evolution of the eastern interior, North America, in Beaumont, C. and Tankard, A. J. eds., Sedimentary basins and basin forming mechanisms, Canadian Society of Petroleum Geologists, Ottawa, v. 12, pp. 425–445. Beuthin, J. D., 1994, A sub-Pennsylvanian paleovalley system in the central Appalachian basin and its implication for tectonic and eustatic control on the origin of the regional Mississippian-Pennsylvanian unconformity, in Dennison, J. M. and Ettensohn, F. R. eds., Tectonic and eustatic controls on sedimentary cycles, SEPM (Society for Sedimentary Geology), Tulsa, OK, v. 4, pp. 107–120. Braille, L. W., Keller, G. R., Hinze, W. J., and Lidiak, E. G., 1982, An ancient rift complex and its relation to contemporary seismicity in the New Madrid Seismic Zone. Tectonics, v. 1, pp. 225–237. Braun, J., and Shaw, R., 2001, A thin-plate model of Palaeozoic deformation of the Australian lithosphere; implications for understanding the dynamics of intracratonic deformation, in Miller, J. A., Holdsworth, R. E., Buick, I. S., and Hand, M. eds., Continental reactivation and reworking, Geological Society of London, London, England, Special Publications, v. 184, pp. 165–193.

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Brodie, J., and White, N., 1995, The link between sedimentary basin inversion and igneous underplating, in Buchanan, J. G. and Buchanan, P. G. eds., Basin inversion, Geological Society of London, Special Publication, v. 88, pp. 21–38. Bunker, B. J., Witzke, B. J., Watney, W. L., and Ludvigson, G. A., 1988, Phanerozoic history of the central midcontinent, United States, in Sloss, L. L. ed., Sedimentary cover — North American craton, U.S., The Geology of North America, Geological Society of America, Boulder, CO, v. D-2, pp. 243–260. Burgess, P. M., and Gurnis, M., 1995, Mechanisms for the formation of cratonic stratigraphic sequences. Earth and Planetary Science Letters, v. 136, pp. 647–663. Burgess, P. M., Gurnis, M., and Moresi, 1997, Formation of North American cratonic sequences by interaction between mantle, eustatic and stratigraphic processes. Bulletin of the Geological Society of America, v. 108, pp. 1515–1535. Burgess, P. M., and Moresi, L. N., 1999, Modelling rates and distribution of subsidence due to dynamic topography over subducting slabs: is it possible to identify dynamic topography from ancient strata? Basin Research, v. 11, pp. 305–314. Burke, K., Macgregor, D. S., and Cameron, N. R., 2003, Africa’s petroleum systems: four tectonic ‘Aces’ in the past 600 million years, in Arthur, T. J., Macgregor, D. S., and Cameron, N. R. eds., Petroleum geology of Africa: new themes and developing technologies, Geological Society, London, Special Publications, v. 207, pp. 21–60. Bushbach, T. C., and Kolata, D. R., 1990, Regional setting of the Illinois Basin, in Leighton, M. W., Kolata, D. R., Oltz, D. F., and Eidel, J. J. eds., Interior Cratonic Basins, Memoir 51, American Association of Petroleum Geologists, Tulsa, OK, pp. 29–55. Catacosinos, P. A., Harrison, W. B., II, and Daniels, P. A., Jr., 1990, Structure, stratigraphy, and petroleum geology of the Michigan Basin, in Leighton, M. W., Kolata, D. R., Oltz, D. 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Kent, D. M., 1987, Paleotectonic controls on sedimentation in the northern Williston basin, Saskatchewan, in Longman, M. W. ed., Williston Basin: anatomy of a cratonic oil province, Rocky Mountain Association of Petroleum Geologists, pp. 45–56. Klein, G. D., 1995, Intracratonic Basins, in Busby, C. J. and Ingersoll, R. V. eds., Tectonics of sedimentary basins, Blackwell Science, pp. 459–478. Kolata, D. R., and Olive, W. W., 1990, Zuni and Tejas sequences Late Cretaceous through Holocene Series, in Leighton, M. W., Kolata, D. R., Oltz, D. F., and Eidel, J. J. eds., Interior cratonic basins, Memoir 51, American Association of Petroleum Geologists, Tulsa, OK, pp. 165–178. Kominz, M. A., and Bond, G. C., 1991, Unusually large subsidence and sea-level events during middle Paleozoic time: new evidence supporting mantle convection models for supercontinent assembly. Geology, v. 19, pp. 56–60. Leighton, M. W., 1990, Introduction to interior cratonic basins, in Leighton, M. W., Kolata, D. R., Oltz, D. F., and Eidel, J. J. eds., Interior cratonic basins, Memoir 51, American Association of Petroleum Geologists, Tulsa, OK, pp. 1–24. Leighton, M. W., and Kolata, D. R., 1990, Selected interior cratonic basins and their place in the scheme of global tectonics, in Leighton, M. W., Kolata, D. R., Oltz, D. F., and Eidel, J. J. eds., Interior cratonic basins, Memoir 51, American Association of Petroleum Geologists, Tulsa, OK, pp. 729–798. Lie, S., and Nummedal, D., 2004, Late Cretaceous subsidence in wyoming: quantiﬁying the dynamic component. Geology, v. 32, pp. 397–400. Marshak, S., Nelson, W. J., and McBride, J. H., 2003, Phanerozoic strike-slip faulting in the continental interior platform of the United States: examples from the Laramide Orogen, Midcontinent, and ancestral Rocky Mountains, in Storti, F., Holdsworth, R. E., and Salvini, F. eds., Intraplate strike-slip deformation belts, Geological Society, London, Special Publications, v. 210, pp. 159–184. Marshak, S., and Paulsen, T., 1996, Midcontinent U.S. fault and fold zones. A legacy of Proterozoic intracratonic extensional tectonism. Geology, v. 24, pp. 151–154. McKenzie, D. P., 1978, Some remarks on the development of sedimentary basins. Earth and Planetary Science Letters, v. 40, pp. 25–32. Mitrovica, J. X., Beaumont, C., and Jarvis, G. T., 1989, Tilting of the continental interiors by the dynamical effects of subduction. Tectonics, v. 8, pp. 1079–1094. Nunn, J. A., Sleep, N. H., and Moore, W. E., 1984, Thermal subsidence and generation of hydrocarbons in Michigan basin. American Association of Petroleum Geologists Bulletin, v. 68, pp. 296–315. Pang, M., and Nummedal, D., 1995, Flexural subsidence and basement tectonics of the Cretaceous Western interior basin, United States. Geology, v. 23, pp. 173–176. Patchett, P. J., Ross, G. M., and Gleason, J. D., 1999, Continental drainage in North America during the Phanerozoic from Nd isotopes. Science, v. 283, pp. 671–673. Patchett, P. J., Embry, A. F., Ross, G. M., Beauchamp, B., Harrison, J. C., Mayr, U., Isachsen, C. E., Rosenburg, E. J., and Spence, G. O., 2004, Sedimentary cover of the Canadian shield through Mesozoic time reﬂected by Nd isotopic and Geochemical results for the Sverdrup Basin, Arctic Canada. The Journal of Geology, v. 112, pp. 39–57. Pysklywec, R. N., and Mitrovica, J. X., 1998, Mantle ﬂow mechanisms for the large scale subsidence of continental interiors. Geology, v. 26, pp. 687–690. Quinlan, G. M., 1987, Models of subsidence mechanisms in intracratonic basins and their applicability to North American examples, in Beaumont, C. and Tankard, A. J. eds., Sedimentary basins and basin-forming mechanisms, Canadian Society of Petroluem Geologists Memoir, v. 12, pp. 463–481. Roksandic, M. M., 1987, The tectonics and evolution of the Hudson Bay region, in Beaumont, C. and Tankard, A. J. eds., Sedimentary basins and basin forming mechanisms, Canadian Society of Petroleum Geologists, Ottawa, v. 12, pp. 507–518. Sanford, B. V., 1987, Paleozoic geology of the hudson platform, in Beaumont, C. and Tankard, A. J. eds., Sedimentary basins and basinforming mechanisms, Canadian Society of Petroleum Geologists, Ottawa, v. 12, pp. 483–505. Sleep, N. H., Nunn, J. A., and Chou, L., 1980, Platform basins. Annual Review of Earth and Planetary Science, v. 8, pp. 17–34. Sloss, L. L., 1963, Sequences in the cratonic interior of North America. Bulletin of the Geological Society of America, v. 74, pp. 93–114. Sloss, L. L., 1988a, Introduction, in Sloss, L. L. ed., Sedimentary cover — North American craton, U.S., Geological Society of America, Boulder, CO, v. D-2, pp. 1–3. Sloss, L. L., ed., 1988b, Tectonic evolution of the craton in Phanerozoic time, in Sedimentary cover — North American craton, U.S. The Geology of North America, Geological Society of America, Boulder, CO, v. 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E. ed., Seismic stratigraphy-applications to hydrocarbon exploration, American Association of Petroleum Geologists Memoir, v. 26, pp. 83–97. Wheeler, P., and White, N., 2000, Quest for dynamic topography: observations from Southeast Asia. Geology, v. 28, pp. 963–966. Ziegler, P. A., 1988, Evolution of the Arctic–North Atlantic and the Western Tethys. American Association of Petroleum Geologists Memoir, v. 89, pp. 198.

CHAPTER 3

Appalachian Foreland Basin of Canada$ Denis Lavoie

Contents 1. 2. 3. 4.

Introduction Regional Geological Setting Tectonostratigraphic Domains of the Appalachians The Taconian-Deformed Basins — The Humber Zone 4.1. End rift–early drift 4.2. The passive margin 4.3. A regional sea-level scenario for the Lower Paleozoic end-rift and passive margin 4.4. The Taconian foreland basin 5. The Post-Taconian to Acadian Basins 5.1. Newfoundland 5.2. Continental Eastern Canada 5.3. Paleogeographic reconstruction of the post-Taconian basins 6. The Sea Level Record in the Lower to Middle Paleozoic Appalachians in Eastern Canada: Eustasy vs. Tectonism 6.1. The Early Cambrian–Late Ordovician Humber Appalachians 6.2. The latest Ordovician to Middle Devonian Acadian basins 7. Hydrocarbon Potential of the Appalachian Basins 7.1. Lower Paleozoic belts — Humber Zone in Que´bec 7.2. Lower Paleozoic belts — Humber Zone in western Newfoundland 7.3. Lower Paleozoic belts — Gaspe´ Belt References

66 66 67 67 69 73 78 78 82 83 84 90 93 93 93 94 94 94 95 95

Abstract The Late Proterozoic breakup of Rodinia led to the formation of Laurentia. The continent had a paleosouthern jagged margin that consisted of recesses and salients; the Canadian segment of that margin belongs to the St. Lawrence Promontory and Quebec Reentrant. The stratigraphic framework and paleogeographic evolution of Cambrian–Ordovician shallow- to deep-marine units deposited during the rift, passive-margin and foreland-basin stages are integrated from western Newfoundland to southern Que´bec. Major sea-level lowstands and highstands are correlated, with some time discrepancy starting to occur in earliest Ordovician. The passive-margin evolution was primarily controlled by eustatic sea-level changes; although some ancestral faults were sporadically active in Late Cambrian to Early Ordovician in the Quebec Reentrant. The diachronous westerly directed late Early to Late Ordovician tectonic-controlled extensional collapse of the shallow-marine foreland shelf from the St. Lawrence Promontory to the Quebec Reentrant was followed by the diachronous collision of volcanic arcs along Laurentia (Taconian Orogeny) which climaxed in the Middle-Late Ordovician interval with collision occurring ﬁrst at the St. Lawrence Promontory. Tectonic quiescence was short-lived along the paleosouthern continental margin of Laurentia as more exotic microcontinents (Ganderia, Avalonia) were closing in. The paleoenvironmental history was signiﬁcantly affected by these Early Silurian to late Early Devonian tectonic events (Salinic and Acadian orogenies). A Late Ordovician to Early Silurian ﬁlling stage was followed by two T-R cycles. The ﬁrst of these cycles was initiated by a tectonically controlled sea-level rise in latest Early Silurian followed by a eustatic sea-level fall in Late Silurian. The cycle culminated in the Salinic unconformity. The second cycle started with a major tectonic collapse in latest Silurian followed by a slow to ultimately rapid sea-level fall from the Early to early Middle Devonian. The cycle ended with sub-aerial exposure and syn-tectonic sedimentation (Acadian Orogeny). $

Geological Survey of Canada Contribution 2002032.

Sedimentary Basins of the World, Volume 5 ISSN 1874-5997, DOI 10.1016/S1874-5997(08)00003-8

r 2008 Elsevier B.V. All rights reserved.

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1. Introduction The recent regional synthesis of the Appalachians in Canada (Williams, 1995) and in the USA (Hatcher et al., 1989) were published as separate volumes, part of the centennial Geological Society of America Decade of North America Geology (DNAG) synthesis endeavor. The information synthesized our knowledge of this mobile belt in the mid- to late-1980s period. These volumes are still a source of invaluable information as they present the most complete and detailed account of the entire Appalachian Orogen. Since publication, critical new research results have become available and allowed a reﬁning of our understanding of this mobile belt. In the Canadian section of the Appalachians, for example, hydrocarbon exploration in the mid-/late-1990s in the Lower and Middle Paleozoic successions (Cooper et al., 2001; Lavoie and Bourque, 2001), a major deep-seismic project (Lithoprobe East, Quinlan, 1998), NATMAP (Canada NATional geoscience MAPping program) regional mapping and thematic studies (Maritimes Basin; 1993–1998, Forelands and Platform; 1999–2004) and two Targeted Geoscience Initiative projects (Red Indian Line; 2000–2003; Appalachian Energy; 2003–2005) prompted some new research activities and resulted in improved understanding of the northern Appalachians (Lynch, 2001; Lavoie et al., 2003a, 2004; van Staal, 2005). This contribution presents an overview of the Lower to Middle Paleozoic stratigraphic architecture, paleogeographic scenarios and relative sea-level history for the evolving sedimentary basins that resulted in the actual northern Appalachians.

2. Regional Geological Setting Rocks ranging from the Neoproterozoic to Cretaceous are found in the Appalachians of North America. This orogenic belt has been shaped by several major tectonic events as well as by local, less severe, but critical tectonic phases (Figure 1). Six signiﬁcant orogenic/deformation events are documented in the Appalachians and are related to the accretion of volcanic arcs, oceanic crust, microcontinents and continents to the progressively more and more composite margin of Laurentia (van Staal et al., 1998; van Staal, 2005; Ettensohn, Chapter 4, this volume). These events are: (1) the Late Cambrian Penobscot Phase (Cambrian oceanic units-Gander) and the coeval Lushs Bight Oceanic Tract-Dashwoods accretion, (2) the end-Middle Ordovician Taconian Orogeny (volcanic arcs-Laurentia), (3) the Silurian Salinic Orogeny (Ganderia-Laurentia), (4) the late Early Devonian Acadian Orogeny (Avalonia-Laurentia), (5) the end-Middle to Late Devonian Neoacadian Orogeny

Figure 1 Tectonic cycles recorded along the ancient continental margin of Laurentia in eastern North America. Major orogenic phases are outlined. The shaded box shows the stratigraphic interval covered in this contribution. Modi¢ed from Sanford (1993).

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(Meguma-Laurentia) and (6) the end-Carboniferous-Permian Alleghanian Orogeny (Gondwana-Laurentia), all these events leading to the formation of Pangea (Miall and Blakey, Chapter 1, this volume). In this contribution, we discuss the ﬁrst four of these events and their tectonostratigraphic effects on the Canadian Appalachian Orogen (Figure 1).

3. Tectonostratigraphic Domains of the Appalachians The tectonostratigraphic domains of the evolving orogenic belt are used to divide the Appalachians into workable packages for geological consideration. The Lower Paleozoic tectonostratigraphic zones (Williams, 1979) include the Humber (Laurentia’s continental domain), Dunnage (peri-Laurentia and peri-Gondwana oceanic domains), the Gander and Avalon zones (peri-Gondwana oceanic and continental domains, respectively) and the Meguma Zone (a late-accreted peri-Gondwana continental terrane) (Figure 2). These belts record the complex evolution of the Cambrian and Ordovician orogenies (Figure 1; van Staal, 2005) and were affected by post-Taconian events that shaped up the Appalachians. The ﬁrst part of this contribution focuses primarily on the Laurentian Humber Zone. The post-Taconian to syn-Acadian basins are developed over the Taconian zones (Figure 2); the best known of these basins is the Gaspe´ Belt that is preserved in various tectonostratigraphic assemblages: the Connecticut Valley–Gaspe´ synclinorium, the Aroostook–Perce´ anticlinorium and the Chaleurs Bay synclinorium. The Early Devonian Acadian Orogeny is the main phase that shaped these elements (Malo and Bourque, 1993; Williams, 1995). The expression of the Silurian Salinic Orogeny (Dunning et al., 1990; Cawood et al., 1994; van Staal, 2005) varies along strike the Appalachians (Waldron et al., 1998; Malo, 2001; Tremblay and Castonguay, 2002; Tremblay and Pinet, 2005; Ettensohn, Chapter 4, this volume).

4. The Taconian-Deformed Basins — The Humber Zone An irregularly shaped continental margin with recesses and salients characterized the southern edge of Laurentia following breakup of Rodinia in Neoproterozoic time (Figure 3; Thomas, 1977, 1991; Miall and Blakey, Chapter 1, this volume). The irregular shape of the margin played a key role in the evolution of the Early

Figure 2 Taconian tectonostratigraphic domains and the Silurian-Devonian basin (Gaspe¤ Belt) of the Canadian Appalachians. The tectonostratigraphic divisions of the Gaspe¤ Belt are shown in Figure 15. Thick red lines refer to the position of seismic lines shown in Figure 4 (southern Quebec), 5 (Western Newfoundland) and 19 (Gaspe¤ Belt). Note the position of the Early Cambrian shallow-marine sediments of the Oak Hill Group in southern Que¤bec. BBL, Baie Verte--Brompton Line; SA, Saguenay Graben. Tectonostratigraphic nappes of the Taconian Humber Zone are detailed in Lavoie et al. (2003b). CB, Silurian Clam Bank Belt. Modi¢ed fromWilliams (1995) and Lavoie et al. (2003b).

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Figure 3 (a) The Early Paleozoic continental margin of Laurentia with the distribution of reentrants and promontories. O-B, Ottawa--BonnecheØre Graben; SA, Saguenay Graben. Modi¢ed from Thomas (1977). (b) General event-stratigraphic framework for the Que¤bec Reentrant (left column) and St. Lawrence Promontory (right column) together with the most signi¢cant tectonic events.

Paleozoic foreland platform in Canada (Stenzel et al., 1990; Lavoie, 1994; Sharma et al., 2003) and in the USA (Quinlan and Beaumont, 1984; Ettensohn, Chapter 4, this volume). Detailed information on the rift, passive-margin and foreland-basin evolution of the shallow-marine Early Paleozoic continental-margin platform is available for western Newfoundland (James et al., 1989) and southern Quebec–eastern Ontario (Bernstein, 1992; Lavoie, 1994, 1995a; Salad Hersi and Dix, 1997; Lavoie and Asselin, 1998; Salad Hersi et al., 2002a, 2003; Salad Hersi and Dix, 2006; Dix and Al Rodhan, 2006). The coeval slope succession has been studied in detail in western Newfoundland (James and Stevens, 1986; James et al., 1989; Waldron and Palmer, 2000; Palmer et al., 2001; Burden et al., 2001; Waldron et al., 2003) and eastern Que´bec (Lebel and Kirkwood, 1998; Lavoie, 1997, 1998, 2001, 2002; Cousineau and Longue´pe´e, 2003; Longue´pe´e and Cousineau, 2005), and a regional integrated framework for this time interval has recently been proposed (Lavoie et al., 2003b). The term Humber Zone (Williams, 1976) was given for the north-westernmost tectonostratigraphic domain of the Taconian orogenic belt (Figure 2). First deﬁned in western Newfoundland, this belt was later recognized and extended on the Canadian mainland down to the northern US segment of the Appalachians (Williams, 1978). In the Humber Zone, stacks of tectonic slices of Neoproterozoic basement and Lower Cambrian to Upper Ordovician rocks of Laurentian continental afﬁnity (St. Lawrence Platform and coeval slope and rise sediments)

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are deformed and thrusted over the St. Lawrence cratonic platform in a thin- to thick-skinned tectonic scenario (St-Julien and Hubert, 1975; Williams, 1978; van Staal et al., 1998; Waldron et al., 1998, 2003; Stockmal et al., 1998; Se´journe´ et al., 2003; Stockmal et al., 2004; Se´journe´ et al., 2005; Figures 4 and 5). The Humber Zone is bordered to the west by the St. Lawrence Platform (Sanford, 1993); the limit is the westernmost transported tectonic slices (Globensky, 1987; Waldron et al., 1998). This limit in southern Que´bec is commonly referred to as the Logan’s line, or as the Champlain Thrust in northern Vermont. Seismic data however, indicate that the St. Lawrence Platform records signiﬁcant Taconian (?) compressive deformation (Figure 4), which include triangle zone and blind thrusts found in the central segment of the St. Lawrence Platform in southern Que´bec (Castonguay et al., 2003, 2006). Therefore in Que´bec, the Appalachian structural front does not coincide with Logan’s line. To the east, the Humber Zone is bordered by the Dunnage Zone, which consists of various oceanic rocks; the western limit of the Humber Zone consists of faults that form the Baie Verte–Brompton Line (Figure 2; Tremblay et al., 1995; van Staal, 2005). The relative timing of obduction of the oceanic seaﬂoor units onto the continental margin in the Que´bec Reentrant has been traditionally indirectly determined by the biostratigraphic age of the successions that underand overlie the accreted units of the Dunnage Zone (St-Julien and Hubert, 1975). Recent 40Ar/39AR and K/Ar metamorphic ages conﬁrm the ‘‘classic’’ Middle to Late Ordovician Taconian age (Castonguay et al., 1997; Pincivy et al., 2003; Glasmacher et al., 2003) for the ophiolite obduction onto the continental slope margin successions. At the St. Lawrence Promontory in western Newfoundland, the obduction of the Bay of Island ophiolite and the associated Humber Arm Allochthon onto the continental margin of Laurentia was long considered to be Middle Ordovician in age (Williams, 1975). This ‘‘classic’’ Taconian age is supported by the biostratigraphic ages of the Taconian ﬂysch and of the overlying units. However, detailed geochronology, structural studies and industry seismic data (Figure 5) indicate that the emplacement of oceanic-domain units over the shallow segment of the continental margin started in Silurian (Dunning et al., 1990; Cawood et al., 1994) and ended prior to the Vise´an (Carboniferous), likely in Middle Devonian (Cawood, 1993; Waldron et al., 1998; Stockmal et al., 1998, 2003, 2004).

4.1. End rift–early drift In western Newfoundland, the oldest rift-related event is dated at 615 Ma (Kamo et al., 1989), however, Cawood et al. (2001) documented that signiﬁcant rifting only started at 570 Ma with a last pulse at 555–550 Ma (van Staal et al., 1998; Waldron and van Staal, 2001; van Staal, 2005). For the western Newfoundland platform succession, James et al. (1989) identiﬁed the end rift–early drift episode as the ‘‘pre-platform shelf ’’ which is recorded by the Lower Cambrian Labrador Group. This event coincides with the Sauk I sub-sequence of the Early Cambrian (Sloss, 1963). In Que´bec, dike-swarm tholeiites in the Grenvillian province give a 590 Ma age (Kamo et al., 1995). Riftrelated alkaline basalts and comendites of the Tibbit Hill Formation in southern Que´bec gave 554 Ma (Kumarapeli et al., 1989) and recent dating of rift-related volcanic rocks in eastern Que´bec yielded ages between 565 +/� 6 Ma and 556 +/� 5 Ma (Hodych and Cox, 2007). There is no unequivocal preserved record of Early Cambrian facies on the St. Lawrence Platform. The Potsdam Group unconformably overlies Precambrian basement; the lower formation (Covey Hill Formation) is assigned an Early Cambrian age (Sanford, 1993) without supporting faunal elements. At the eastern end of the Humber Zone (Figure 2), tectonic stacks of the shallow-marine Oak Hill Group (Charbonneau, 1980) overlie rift volcanics (Kumarapeli et al., 1989; Castonguay et al., 2001). The Cheshire (sandstone) and Dunham (carbonate) formations have yielded Early Cambrian faunal elements (Clark, 1936; Clark and McGerrigle, 1944).

4.1.1. The St. Lawrence promontory (western Newfoundland) The Lower Paleozoic continental margin of Laurentia preserved in western Newfoundland was built on the St. Lawrence Promontory (Figure 3) (Thomas, 1977, 1991; Miall and Blakey, Chapter 1, this volume). Autochthonous and transported rocks are preserved. The best known of the allochthons is the Humber Arm Allochthon (Williams, 1975). Its sedimentary rock package forms the Humber Arm Supergroup (Stevens, 1970). The current framework includes a Neoproterozoic (?)-Lower Cambrian Curling Group (Summerside, Irishtown and Blow Me Down Brook formations), a Middle Cambrian-mid-Arenigian Northern Head Group (Cooks Brook and Middle Arm Point formations) and the coeval, laterally correlative Cow Head Group (Shallow Bay and Green Point formations) (Figure 6). The lower assemblage (the Curling Group) has no equivalent below the Cow Head Group (Figure 6). The slope and rise sediments of the end-rift episode are recorded in the Curling Group (Lavoie et al., 2003b).

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Figure 4 Seismic lines M2001 and M2002 in southern Quebec imaging the St. Lawrence Platform, the Appalachian structural front and the stack of nappes in the Humber domain. Only the northwestern segment of line M2001 is shown. Note the development of signi¢cant compressive structures west of the Logan’s line, in particular, a shallow triangle zone on M2002. The position of the Saint-Flavien gas ¢eld is shown on the M2001 line. The deformation is thin-skinned. Modi¢ed from Castonguay et al. (2006). Location of the lines in southern Quebec is shown by the red lines in Figure 2.

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Figure 5 Seismic line 91-20 o¡shore western Newfoundland imaging the platform and foreland-basin successions involved in thick-skinned deformation with basin inversion along the Round Head Thrust Fault (RHT) and the development of a triangle zone. The projected position of the onshore Garden Hill (GH) oil ¢eld is shown. Modi¢ed from Stockmal et al. (1998). RIT, Red Island Thrust; SGT, St. George Thrust; TCT,Tea Cove Thrust. Location of the line is shown by the blue line in Figure 2.

The Neoproterozoic to Lower Cambrian Summerside Formation consists of slates with subordinate metasandstones and conglomerates (Stevens, 1965, 1970; Waldron and Palmer, 2000; Palmer et al., 2001; Waldron et al., 2003). The overlying Lower Cambrian Irishtown Formation consists of slates with sandstones and limestone conglomerates. (Palmer et al., 2001). The upper Lower Cambrian Blow Me Down Brook Formation (Botsford, 1988; Lindholm and Casey, 1990; Burden et al., 2001) consists of parallel and cross-laminated, quartzrich feldspathic sandstone with shale (Waldron and Palmer, 2000; Buchanan et al., 2001; Waldron et al., 2001, 2003). The Curling Group is time-correlative with the Lower Cambrian shallow-marine Labrador Group (James et al., 1989; Figure 6). Microfaunal correlations have been proposed between the Curling Group and the Forteau Formation (Labrador Group) (Burden et al., 2001; Normore, 2001). The upper unit of the Labrador Group, the Hawke Bay Formation, records a major sea-level lowstand (James et al., 1989). It has been proposed that massive sandstone and conglomerate in the upper part of the Irishtown Formation represents the slope record of that major lowstand (James et al., 1989; Palmer et al., 2001; Lavoie et al., 2003b).

4.1.2. The Que´bec reentrant (eastern Que´bec) In eastern Que´bec, the Lower Paleozoic continental margin of Laurentia was built in the Que´bec Reentrant (Figure 3) (Thomas, 1977, 1991; Miall and Blakey, Chapter 1, this volume). The Humber succession in the Que´bec Reentrant occurs in a number of stacked structural nappes (Figure 2), for which stratigraphic nomenclatures were only recently synthesized (Lavoie et al., 2003b) (Figures 2 and 7). At the base of the Humber succession, the undated Saint-Roch Group (and correlative units; Figure 7) consists of mudstone with subordinate sandstone and rift volcanics (Lavoie, 1997). A distinctive unit of massive, pebbly green sandstone with red and green mudstone of late Early Cambrian age (Saint-Nicolas — ‘‘green

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Figure 6 (a) End-rift (green colored) and passive-margin (blue colored) stratigraphic correlation of platform succession and Humber Zone continental-slope sediments in western Newfoundland. The Humber Arm Supergroup consists of the lower Curling Group and the laterally equivalent Cow Head and Northern Head groups. Designation of units as follows: GROUP, Formation, Member. Foreland-basin units are in italic. Step, Steptoan; Sunw, Sunwaptan; Skull, Skullrockian; Trem., Tremadocian; Darr., Darriwilian. (b) General lateral relationship of stratigraphic units from platform to distal slope. BMDB, Blow Me Down Brook Formation; Summer., Summerside Formation. Red-colored intervals represent a hiatus. Not to scale. Modi¢ed from James et al. (1989) and Lavoie et al. (2003b).

sandstone’’ and correlative units; Figure 7) overlies the basal succession (Sweet and Narbonne, 1993; Lavoie et al., 2003b; Burden, 2003). This distinctive massive sandstone unit is a regional marker. It has been proposed that this coarse-grained unit represents the deep-marine expression of a signiﬁcant late Early Cambrian sea-level lowstand (Lavoie et al., 2003b). 4.1.3. Correlation western Newfoundland–Que´bec Following the Late Neoproterozoic initiation of spreading and eruption of basalts with associated coarse-grained sedimentation (Wood Island Lavas, Bradore, Saint-Anselme and Tibbit Hill formations, Caldwell and Shickshock

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Figure 7 End-rift (green colored) and passive-margin (blue colored) stratigraphic correlations of platform and Humber Zone continental-slope sediments in the Que¤bec Reentrant in (a) southern Que¤bec and (b) eastern Que¤bec-Gaspe¤. See Figure 6 for unit status. Autoch., Autochthonous units of the St. Lawrence Platform; Alloch., Allochthonous platform thrusts in the Humber Zone; BEEKMAN., Beekmantown Group; Hasting Cr., Hasting Creek Formation; Naylor Led., Naylor Ledge Formation; Kam, Kamouraska Formation; L.C., Le¤vis Formation conglomerate. Red-colored intervals are for hiatus. (c) Schematic composite diagram of the St. Lawrence Platform and Humber Zone illustrating the assumed palinspastically restored lateral relationships of some tectonostratigraphic nappes. Note the thickening of the sedimentary pile toward the east. The Early Cambrian platform units of the Oak Hill Group are part of a distinct tectonic nappe.Vertical exaggeration 15x. Modi¢ed from Lebel and Kirkwood (1998).

groups), the ensuing relative sea-level rise led to shallow-marine carbonate-siliciclastic sedimentation on local horst structures (Forteau and Dunham formations), whereas in the graben and slope settings, ﬁne- and coarsegrained sediments were deposited as proximal and distal submarine fans (Summerside, Irishtown, Blow Me Down Brook, Sainte-Foy and Armagh formations, lower beds of Saint-Roch Group) (Figures 8 and 9a). A major sealevel lowstand is recognized in late Early Cambrian and marks the end of the Sauk I sub-sequence (James et al., 1989; Lavoie et al., 2003b). This lowstand is recorded in prograding shallow-marine sandstone such as the Hawke Bay (Newfoundland; Knight and Boyce, 1987) and Monkton (Vermont; Landing et al., 2002) formations. This event is expressed in the deep-marine sandstone and conglomerate found in the upper part of the Irishtown, in the Blow Me Down Brook and Saint-Nicolas formations and the ‘‘green sandstone’’ unit of the Saint-Roch Group (Figures 8 and 9b).

4.2. The passive margin In western Newfoundland, the late Early Cambrian shallow-marine clastics were ﬂooded by the transgressive sea level that marks the initiation of the Sauk II sub-sequence (James et al., 1989). The carbonate-dominated passive margin consisted in a Middle-Late Cambrian narrow, high-energy carbonate platform (Port au Port Group) that evolved into an Early-earliest Middle Ordovician wide, low-energy carbonate platform (St. George Group) (Figure 6; James et al., 1989). The passive-margin period ended with onset of signiﬁcant seaﬂoor subduction, and

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Figure 8 Stratigraphic correlation of the end-rift (Sauk I sub-sequence; green colored) and passive-margin (Sauk II and III; blue colored) successions in the Que¤bec Reentrant and at the St. Lawrence Promontory. Correlations for platform and slope rock units are proposed; see text for details. Red colored intervals are for hiatus. See Figures 6 and 7 for unit status and acronyms. Modi¢ed from Lavoie et al. (2003b).

the migration of a tectonic peripheral bulge on the margin in earliest Middle Ordovician (Jacobi, 1981; Knight et al., 1991). This event coincides with the end of the Sauk sequence, the beginning of the Tippecanoe sequence and the inception of the Taconian foreland basin (Figures 1 and 3). In the Que´bec Reentrant (Figure 7), the oldest known passive-margin, shallow-marine platforms are the upper Middle Cambrian Corner-of-the-Beach Formation (Kindle, 1942; Lavoie, 2001) and the shallow-marine carbonate platform of the Upper Cambrian Strites Pond Formation (Salad Hersi and Lavoie, 2001a; Salad Hersi et al., 2002b). These two formations are facies-wise similar to the Port au Port Group (Lavoie, 2001; Salad Hersi and Lavoie, 2001a; Salad Hersi et al., 2002b). The shallow-marine record of the Sauk II and III sub-sequences is best expressed in the extensive Lower Ordovician carbonates of the Beekmantown and Philipsburg groups of southern Que´bec (Globensky, 1987; Bernstein, 1992; Salad Hersi et al., 2002a, 2002b, 2003) as well as of the Romaine Formation on Mingan Islands (Desrochers, 1988). The Lower Ordovician units are truncated by the unconformity resulting from the migration of the peripheral bulge (Knight et al., 1991; Lavoie, 1994).

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Figure 9 (a) Interpreted paleogeographic reconstruction of the Laurentian continental margin in Early Cambrian. Geographically restricted carbonate platforms (Forteau/Dunham) are located on rift-related horsts, while intervening deep grabens are ¢lled by coarse- to ¢ne-grained slope sediments (Armagh/Saint-Roch/Summerside). The thick dashed line represents the assumed position of the continental shelf edge. (b) Paleogeographic reconstruction in late Early Cambrian at the time of the ‘‘Hawke Bay event’’ (¢rst major sea level lowstand) that marks the end-rift episode. Shallow-marine quartzite (Monkton/Hawke Bay) prograded toward the shelf break while deeper slope environments were fed by coarse-grained sediments (‘‘Green Sandstone’’/Irishtown) from the prograding clastic wedge. Modi¢ed from Lavoie et al. (2003b). The cartoon illustrates the late Early Cambrian paleoenvironmental model; the progradation of the clastic wedge (Hawke Bay Formation) responds to the major sea-level lowstand. Cartoon modi¢ed from James et al. (1989).

4.2.1. The St. Lawrence promontory (western Newfoundland) The slope record of the passive margin consists of two laterally correlative rock packages: (1) the Middle Cambrian-lowermost Middle Ordovician proximal Cow Head Group and (2) the coeval, but more distal, succession of the Northern Head Group (Figure 6). These deep-marine successions are well dated; limestone fragments in the conglomerates that punctuate the succession are rich in shallow-marine fauna, whereas the intervening ﬁne-grained sediments are dated by graptolites and acritarchs. The Cow Head Group has been extensively studied (James and Stevens, 1986; Coniglio, 1986; James et al., 1989) and consists of two laterally correlative formations, the proximal Shallow Bay and the distal Green Point formations (Figure 6). Seven distinctive rock assemblages (deﬁned as members) of alternating shales, sandstones and ﬁne- to coarse-grained limestones are recognized in the Shallow Bay (four members) and Green Point (three members) formations (Figure 6). The biostratigraphy allows correlation between both formations, with the shallow-marine platform and with the Cambrian Grand Cycles framework (James and Stevens, 1986; James et al., 1989).

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The Northern Head Group (Botsford, 1988) consists of the Cooks Brook and Middle Arm Point formations (Figure 6). The limestone and shale succession of the upper Middle Cambrian to Lower Ordovician Cooks Brook Formation disconformably overlies the upper Lower Cambrian Irishtown Formation (Figure 6). The Middle Arm Point Formation consists of mudstone with subordinate silty dolostone and limestone that carry Tremadocian to Arenigian age graptolite. The passive-margin succession in western Newfoundland consists of proximal (Cow Head Group) and distal (Northern Head Group) depositional assemblages (Figure 6). The detailed work by James and Stevens (1986) and James et al. (1989) correlates the evolution of the Cow Head Group with that of the coeval shallow-marine Port au Port and St. George groups (Figure 6). Two major T-R cycles (Sauk II and III sub-sequences) are recorded in the Cow Head Group. Correlation of the Cow Head Group with the Northern Head Group is also proposed (Figure 6) (Lavoie et al., 2003b). 4.2.2. The Que´bec reentrant (eastern Que´bec) The ﬁrst passive-margin sediments that overlie the upper Lower Cambrian green sandstone unit consist of a thick succession of upper Lower Cambrian to lower Middle Cambrian mudstone with glauconite- and quartz-rich sandstone (Orignal Formation and correlative units; Figure 7). A distinctive coarse-grained unit (St-Damase Formation and correlative units; Figure 7) overlies this ﬁne-grained dominated interval. This unit consists of channel-ﬁll carbonate conglomerate, feldspathic and siliceous sandstone and minor mudstone (Lavoie, 1998). The matrix of the conglomerate is Late Cambrian and embedded fragments consist of Early to early Late Cambrian nearshore to shallow-marine platform-margin limestone facies together with meter-sized sandstone, basalt fragments and basement-derived gneiss and orthoquartzite (Lavoie, 1997, 1998). The coarse-grained interval is overlain by a succession of uppermost Cambrian to lowermost Ordovician mudstone with subordinate sandstone (Rivie`re-du-Loup Formation and correlative units; Figure 7) with discontinuous thick channel-ﬁll quartz arenite (Kamouraska Formation; Figure 7). The youngest passive-margin unit (Rivie`re Ouelle Formation and correlative units; Figure 7) consists of variegated mudstone with subordinate sandstone, ribbon limestone, calcarenite and limestone conglomerate of Early Ordovician age (Landing and Benus, 1985; Landing et al., 1986; Bernstein et al., 1992; Maletz, 1992, 2001; Asselin and Achab, 2004). 4.2.3. Correlation Newfoundland–Que´bec The passive-margin history consists of two major T-R cycles identiﬁed as the Sauk II and III sub-sequences (Figure 8). The initial Sauk II transgressive and early highstand sea level (Middle Cambrian) resulted in the sedimentation of the lower part of the Port au Port Group with no preserved slope record until the late Middle Cambrian (Cooks Brook Formation). In eastern Que´bec, the Corner-of-the-Beach Formation indicates that the Middle Cambrian platform locally extended into the more outer part of the Que´bec Reentrant; the slope record consists only of mudstone and sandstone (e.g., Orignal Formation and equivalent units). A sea-level lowstand is recognized near the base of the Late Cambrian (Steptoan) and correlates with the end of Grand Cycle A (Chow and James, 1987; Cowan and James, 1993), whereas the slope record of the marine lowstand is found in limestone conglomerates at the base of the Shallow Bay and Cooks Brook formations. These conglomerates consist of limestone fragments derived from late highstand to lowstand shedding of platform-margin facies (James et al., 1989). In Que´bec, no record of lower Upper Cambrian platform rocks is known, but the widespread lower Upper Cambrian (Steptoan) slope limestone conglomerates (Saint-Damase Formation and correlative units) indicate the former presence of this platform. These thick conglomerates correlate with the end of the Sauk II sub-sequence. Presence of the pre-Upper Cambrian fragments in the conglomerate makes the Upper Cambrian conglomerates of the Que´bec Reentrant different from the Newfoundland time correlatives. Simple late highstand to lowstand shedding of platform-margin clasts cannot alone explain the presence of fragments of Neoproterozoic (basement) to Late Cambrian age. The thickest and best-developed conglomerate successions are found in eastern Quebec, adjacent to the Saguenay Graben (Aulacogene of Kumarapeli and Saull, 1966; Figures 2 and 10a) and a Late Cambrian reactivation of this failed Neoproterozoic rift has been proposed as causal mechanism (Lavoie et al., 2003b). This erosion of the continental margin ended with the ﬁnal exhumation of basement rocks through erosion of the carbonate, clastic and volcanic succession deposited during and after the rift episode. The exact cause for the reactivation of the Saguenay Graben is still unclear; however, the ﬁrst accretionary events recorded in circum-Iapetus terranes occurred in late Middle Cambrian (Miall and Blakey, Chapter 1, this volume). This is well-documented by the Penobscotian Phase in the units of the peri-Gondwana Gander terrane (Neuman, 1967) and in the obduction of the peri-Laurentia Bushs Bight Oceanic Tract over the

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Figure 10 (a) Interpreted paleogeographic reconstruction of the Laurentian continental margin in Late Cambrian (Skullrockian). Lateral relationship between the Cairnside and Strites Pond formations is discussed in Salad Hersi et al. (2002a, 2002b).The thickest succession of the Upper Cambrian conglomerates (Saint-Damase and correlative units) with older basement and platform rocks are found in the Lower St. Lawrence Valley, near the Saguenay Graben (SA).The two cartoons illustrate the depositional settings and lateral platform --- slope relationships in Middle-Late Cambrian. Not to scale. (b) Interpreted paleogeographic reconstruction of the Laurentian continental margin in Early Ordovician (Arenigian).The lateral relationship between the Theresa and Wallace Creek formations is discussed in Salad Hersi et al. (2002b, 2003).Thick white dashed lines represent the hypothetical facies transition from proximal to distal slope facies.The diachronism in Early Ordovician sea level lowstands (Tremadocian-Arenigian inWestern Newfoundland and mid-Arenigian in southern Que¤bec) is interpreted to have been related to the tectonic in£uence of the Ottawa--BonnecheØre Graben (O-B).The two cartoons illustrate the depositional settings and lateral platform--slope relationships in Early Ordovician. Not to scale.

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Dashwoods microcontinent (van Staal et al., 2004). The latter accretion occurred in the neighborhood of Laurentia and could have been the trigger for Late Cambrian tectonic instability in the Quebec Reentrant. At the onset of the Sauk III sub-sequence (Figure 8), a sea-level highstand is recorded in the Grand Cycle C and is dated as Late Steptoan. On the Newfoundland platform, this marine highstand is recorded in the Berry Head Formation (Port au Port Group); in the slope environment, the ﬁner-grained units of the Cow Head Formation record this high sea level. No carbonate platform of that age is known in the Que´bec Reentrant, but a ﬁne-grained siliciclastic slope record (Rivie`re-du-Loup Formation and equivalents units) likely correlates with the Newfoundland slope succession. This highstand was followed by another sea-level lowstand in latest Cambrian (Early Skullrockian), which marks the end of the Port au Port Group in western Newfoundland. In the slope environment, the ﬁne-grained succession of the Cow Head Group passes into a more conglomeratic section. In southern Que´bec, the platform carbonates of the Strites Pond Formation are dated as latest Cambrian (Skullrockian), and is capped by an erosive sub-aerial unconformity, which separates it from the outer-shelf facies of the upper Tremadocian Wallace Creek Formation (Salad Hersi and Lavoie, 2001a; Salad Hersi et al., 2002b). The coeval slope record of this erosive event is the thick channel-ﬁll quartz sands of the Kamouraska Formation. The Early Ordovician was marked by craton-wide transgression leading to the deposition of the St. George Group in western Newfoundland (Pratt and James, 1986), the Romaine Formation on Mingan and Anticosti islands (Desrochers, 1988; Brennan-Alpert, 2001) and the Beekmantown Group in southern Que´bec and eastern Ontario (Bernstein, 1992; Salad Hersi et al., 2002a, 2003) (Figure 8). Two T-R depositional cycles are recognized; the mid-Boat Harbour Formation unconformity (Tremadocian-Arenigian boundary) separates these two cycles (James et al., 1989) as does the slightly younger (mid-Arenigian) Theresa–Beauharnois formation contact (Salad Hersi et al., 2003; Dix and Salad Hersi, 2004; Dix and Al Rodhan, 2006). In western Newfoundland, the ﬁrst cycle covers the Tremadocian; the second cycle is Arenigian to Darriwilian in age. In southern Quebec and eastern Ontario, the T-R cycles are seemingly younger; the ﬁrst cycle is upper Tremadocian to mid-Arenigian whereas the second one is mid-Arenigian to Darriwilian in age (Dix and Salad Hersi, 2004). Tectonic instability associated with the Ottawa–Bonneche`re Graben has been proposed to explain the diachronism in the sea-level ﬂuctuations, conversely imprecise biostratigraphic data could also be envisaged (Dix and Salad Hersi, 2004; Dix and Al Rodhan, 2006). In the coeval slope succession, these limits of the T-R cycles are expressed in shedding of major limestone conglomerates in the Shallow Bay (James et al., 1989) or in the Le´vis (Samson et al., 2002; Lavoie and Kirkwood, 2006) formations (Figure 8).

4.3. A regional sea-level scenario for the Lower Paleozoic end-rift and passive margin The proposed Newfoundland–Quebec correlation allows the recognition of four distinctive major sea-level lowstands (Figure 11): (1) A late Early Cambrian event (‘‘Hawke Bay event’’) expressed by massive sandstones and conglomerates. This event marks the upper limit of the Sauk I sub-sequence. (2) An early-mid Upper Cambrian event (Steptoan-Sunwaptan) expressed in limestone conglomerates on the slope. Coeval tectonic instability is indirectly documented near the Saguenay Graben in the Que´bec Reentrant. This event marks the end of the Sauk II sub-sequence. (3) A latest Cambrian event (Early Skullrockian) represented by local limestone conglomerates and coarsegrained clastics. This event occurs within the Sauk III sub-sequence but marks the end of the Cambrian Grand Cycle C. (4) (a) In western Newfoundland, a sea-level lowstand at the Tremadocian–Arenigian boundary represented by more limestone conglomerates and (b) in southern Que´bec–eastern Ontario, a slightly younger sea-level lowstand in mid-Arenigian expressed in the major limestone conglomerate of the Le´vis Formation. Diachronism is related to tectonic instability of the Ottawa–Bonneche`re Graben (Dix and Salad Hersi, 2004). The stratigraphic record indicates that tectonism was sporadically active in the Quebec Reentrant from Late Cambrian to Early Ordovician (Lavoie et al., 2001a; Lavoie et al., 2003b; Dix and Salad Hersi, 2004; Dix and Al Rodhan, 2006). Evidence for such instability is associated with the two failed aulacogens in the Quebec Reentrant, the Ottawa–Bonneche`re and the Saguenay grabens.

4.4. The Taconian foreland basin The building of passive-margin successions along the eastern seaboard of Laurentia was stopped by emergence and sub-aerial exposure of the platform in earliest Middle Ordovician. The resulting break is known variously as

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Figure 11 Chronostratigraphic correlation for selected units of the end-rift--passive-margin successions along the entire Canadian segment of Laurentia. The four major sea-level lowstands discussed in text (blue-colored intervals) are identi¢ed (1--3 and the slightly diachronous 4a and 4b). Depositional hiatus are in red.The sea-level curve is from James et al. (1989); the curve applies well to the Que¤bec succession prior to the Early Ordovician. The dashed line displays the Quebec Reentrant departure from the James et al. (1989) curve. Correlation with Cambrian grand cycles is shown. MH PT, March Point Fm.; Berry HD, Berry Head Fm.; Agua, Aguathuna Fm.; Summ, Summerside Fm.; Irish, Irishtown Fm.; Covey H, Covey Hill Fm.; Strites P, Strites Pond Fm.; Wal Cr,Wallace Creek Fm.; MC, Morgan’s Corner Fm.; Hast Cr, Hasting Creek Fm.; NL, Naylor Ledge Fm.; St.Dama, Saint-Damase Fm.; RDL, RivieØre-du-Loup Fm.; Kamour, Kamouraska Fm.; L.C., Le¤vis Formation. Modi¢ed from Lavoie et al. (2003b).

the St. George (Newfoundland; Knight et al., 1991), the Romaine (Anticosti; Desrochers, 1988), the Beekmantown (southern Que´bec; Dykstra and Longman, 1995; Salad Hersi et al., 2003), the intra-Philipsburg (southern Que´bec–Vermont; Knight et al., 1991), or the Knox (east USA; Read, 1989; Ettensohn, Chapter 4, this volume) unconformity. This unconformity coincides with the boundary between Sloss’ (1963) Sauk and Tippecanoe Sequences and marks the inception of the foreland basin at Laurentian continental margin (Figures 1 and 3). The evolution of the foreland basin became strongly diachronous along strike of the continental margin, which suggests that the reentrant-promontory morphology played a key role in the evolution of Laurentia (Lavoie, 1994; Ettensohn, Chapter 4, this volume).

4.4.1. The St. Lawrence promontory (western Newfoundland) At the St. Lawrence Promontory (Figure 3), the migration of a tectonic peripheral bulge led to compression, block faulting, uplift and erosion of the St. George carbonate platform. The following subsidence contributed to mid-Darriwilian carbonate sedimentation (Table Point Formation; Stenzel et al., 1990) (Figure 12). Continued subsidence led to deep-marine carbonate-shale and eventually to deep-marine shales. Early tectonic exhumation

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Figure 12 Middle to Late Ordovician foreland-shelf to slope stratigraphic framework at the St. Lawrence Promontory and in the Quebec Reentrant. In Newfoundland, the shallow-marine units are restricted to the Darriwilian and the shelf was buried by westerly migrating Taconian £ysch (green colored) by Late Darriwilian. CC, Cap Cormorant Formation; TC,Table Cove Formation; BC, Black Cove Formation. In the Quebec Reentrant, a progressive westward foundering (green-colored £ysch units) of shallow-marine settings occurred in Late Ordovician (Caradocian--Ashgillian). Taconian £ysch sedimentation on the slope was initiated in Late Arenigian (Newfoundland) to early Darriwilian time. Red-colored intervals are for depositional hiatus. The sea-level curve of Ross and Ross (1988) indicates that the evolution of the Newfoundland succession was controlled primarily by tectonism. The Que¤bec Reentrant successions locally record eustatic events (lowstand in yellow, highstands in blue). F, relative sea-level fall; R, relative sea-level rise.

along the Round Head Precursor Fault (Waldron et al., 1993; Stockmal et al., 1998) resulted in local submarine erosion of tectonic escarpments and sedimentation of fault-scarp conglomerates (Cap Cormorant Formation). In Late Darriwilian, the Taconian-derived foreland ﬂysch (Mainland Sandstone of the Goose Tickle Group; Quinn, 1995) was deposited on the foundered platform (Figure 12). The transition from the passive margin to a foreland basin is well expressed in the slope and rise environment (Figure 12). Greenish ﬂyschoid sandstone with subordinate shale of the Middle to Upper Darriwilian Lower Head Formation (James and Stevens, 1986) conformably overlies the proximal carbonate-rich succession of the Cow Head Group. Flyschoid sandstone with shale of the Arenigian to Darriwilian Eagle Island Formation (Waldron et al., 1998; Waldron and Palmer, 2000) overlies the Middle Arm Point Formation of the Northern Head Group. Taconian ﬂysch are demonstrably slightly older in the more distal successions (Figure 12). 4.4.2. The Que´bec reentrant (eastern Que´bec) The inception of the foreland basin and increased tectonic subsidence in the Que´bec Reentrant (Figure 3) was marked by a signiﬁcant change in the St. Lawrence platform (Sanford, 1993). Siliciclastic units covered the

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unconformity (Globensky, 1987; Desrochers, 1988; Salad Hersi and Dix, 1997), and were followed by a thick succession of ramp carbonates in a tectonically active environment (Desrochers, 1988; Sanford, 1993; Lavoie, 1994, 1995a; Sharma et al., 2003; Lemieux et al., 2003). The carbonate sedimentation was shut down diachronously in a westerly direction, with the sedimentation of deep-marine shales and overlying Taconian ﬂysch and ﬁnal molasse (Sanford, 1993; Lavoie, 1994; Sharma et al., 2003) (Figure 12). A similar westward ‘‘younging’’ of ﬂysch sedimentation is also noted for coeval successions in eastern USA (Ettensohn, Chapter 4, this volume). In the Que´bec Reentrant, the carbonate sedimentation lasted from Darriwilian to Late Caradocian. The passive-margin slope deposits are overlain by Darriwilian to lowermost Caradocian ﬂyschoid sandstone with subordinate mudstone, calcarenite, conglomerate and chert (Tourelle Formation and equivalent units, Figure 12) (Clark and Globensky, 1973; Biron, 1974; Hiscott, 1978; De Broucker, 1986; Slivitzky and St-Julien, 1987; Slivitzky et al., 1991; Bloechl, 1996; Prave et al., 2000). Peculiar me´langes are widely distributed and are interpreted to be roughly coeval with these Middle Ordovician units (Figure 12). The best exposed of these me´langes is the Cap Chat Me´lange (Cousineau, 1998). This me´lange consists of broken units of the adjacent formations, in particular, cm- to km-sized blocks of the Rivie`re Ouelle, Tourelle and Des Landes formations (Arenigian to Darriwilian) in a muddy to sandy matrix. In southern Que´bec (Figure 12), chaotic units described as polymictic conglomerate (Citadelle Formation; Osborne, 1956), olistostrome (Drummondville Olistostrome; Slivitzky and St-Julien, 1987) and tectonosome (Pointe-Aubin me´lange; Comeau et al., 2004) are exposed. These chaotic units in southern Que´bec differ from the Cap Chat Me´lange; they are composed of small- to large-sized blocks of the various lithologies found in the shallow- and deep-marine passive-margin and foreland-basin succession. Syn-orogenic sedimentation lasted until end-Caradocian in the Gaspe´ Peninsula and is represented by Upper Ordovician coarse and ﬁne-grained ﬂysch (Cloridorme Formation; Enos, 1969; Prave et al., 2000) (Figure 12). The Cloridorme Formation is coeval with ﬂysch on the St. Lawrence Platform (Globensky, 1987) (Figure 12).

4.4.3. Correlation Newfoundland–Que´bec The evolution of the end-rift and passive-margin episodes was primarily controlled by eustasy with tectonism locally recognized in the proximity of failed rift grabens (Ottawa–Bonneche`re and Saguenay grabens). Differences are recorded at the onset of closure of the Humber Seaway (Waldron and van Staal, 2001), a sea arm of the Iapetus Ocean that separated the Dashwoods microcontinent from Laurentia. These differences are expressed in the timing of various sedimentary events (Figure 12). This diachronous evolution suggests that the overriding control on development and evolution of depositional successions was tectonic (see also Ettensohn, Chapter 4, this volume). The Middle to Late Ordovician eustatic sea-level curve of Ross and Ross (1988) suggests three transgressive–regressive (T-R) events in that period (Darriwilian to end-Ashgillian) corresponding to Sloss’s (1963) Tippecanoe I Sub-sequence (Figure 12). These T-R cycles are: (1) Darriwilian, (2) Caradocian and (3) Ashgillian (Figure 12). Glacio-eustatic processes controlled the last one of these cycles (Brenchley et al., 1994; Gibbs et al., 1997; Lavoie and Asselin, 1998). The foundering of the continental margin at the St. Lawrence Promontory occurred in the Middle to Late Darriwilian (Stenzel et al., 1990; Quinn, 1995) at a time of eustatic sea-level fall during the ﬁrst T-R cycle (Figure 12); this suggests an overriding tectonic control on facies architecture and evolution. The coeval record in the Que´bec Reentrant suggests the presence of an overall transgressive–regressive cycle, which starts at the SaukTippecanoe unconformity and ends in an unconformity that separates the Chazy and Black River groups (Salad Hersi and Lavoie, 2001b; Dix, 2003) (Figure 12). This suggests that the Darriwilian succession in the Que´becOntario shallow-foreland recorded a eustatic signal. The Caradocian and Ashgillian T-R eustatic cycles are imperfectly recorded; the transition from carbonate ramp (Trenton Group) to deep-marine sediments (Utica) and ﬂysch (Sainte-Rosalie) records tectonically driven deeper marine conditions (Lavoie, 1994; Lavoie and Asselin, 1998) (Figure 12). Detailed sedimentologic analyses are unavailable for the deeper marine successions preserved in the Taconian allochthons. The lateral variation in time and nature of middle Upper Ordovician Me´langes in the Que´bec Reentrant and St. Lawrence Promontory reﬂects foreland-propagating compression and local exhumation and erosion of different segments of the foreland and older passive-margin facies. The propagation of compressive deformation and stacking can form depositional basins fed by the exposed succession at the top of the structural stack (Stockmal et al., 2003). The mid-Darriwilian Cape Cormorant in western Newfoundland represents one of these proximal tectonic basins (Stenzel et al., 1990; Waldron et al., 1993). In the Quebec Reentrant, the Darriwilian Cap Chat Me´lange formed far away from the shelf break as it only consists of cannibalized coeval to slightly older deep-marine units. The Darriwilian chaotic units in southern Que´bec formed closer to the shelf break as suggested by the nature of the fragments (Comeau et al., 2004). These basins reached the shelf in

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Caradocian resulting in the Lacolle Breccia, for which event and facies correlations with the Darriwilian Cape Cormorant Conglomerate has been proposed (Lavoie, 1994).

5. The Post-Taconian to Acadian Basins In eastern North America, the timing and physical expression of the Taconian Orogeny varies signiﬁcantly (this contribution and Ettensohn, Chapter 4, this volume). Along strike in the Appalachians, signiﬁcant stratigraphic differences are known in the Upper Ordovician to Middle Devonian successions, which led previous workers to propose distinct sedimentary basins. These basins may overlie more than one of the previous Taconian zones and are spatially distinguished and deﬁned on lithologic, paleontological and structural elements. These basins overlie Taconian zones either unconformably or paraconformably. The most pronounced unconformities occur where the Middle Paleozoic successions overlie the Humber Zone. In most cases, where Middle Paleozoic strata overlie the Dunnage Zone, the contact is a paraconformity with a more or less important time hiatus. The inception of post-Taconian successions follows the westward diachronous sedimentation of ﬂysch. The oldest post-ﬂysch strata (end-Darriwilian to Caradocian) are found at the St. Lawrence Promontory or close to it (in southern Gaspe´ and Te´miscouata) (Figure 13). In the most inner segments of the Que´bec Reentrant

Figure 13 Stratigraphic relationships between the Taconian foreland-shelf and £ysch and the overlying post-Taconian basins. Note that the time hiatus (red-colored intervals) increases from Newfoundland to northern Gaspe¤. The post-Taconian £ysch succession of the Clam Bank Belt in western Newfoundland is detailed. The base of the Long Point Group (Lourdes,Winterhouse and Misty Point formations) is older than the end of Taconian foreland-shelf and £ysch in the Que¤bec Reentrant. The Ross and Ross (1988) sea-level curve for the Late Ordovician indicates a Caradocian T-R cycle correlative with the one recorded by the Long Point Group. The Lower Devonian Clam Bank Formation unconformably overlies the Long Point Group with no preserved record Silurian strata. The Dennison (1985) eustatic sea-level curve for the Early Devonian suggests initial transgressive conditions followed by relative sea-level fall; the Clam Bank Formation records this eustatic signal.

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(Connecticut Valley–Gaspe´ synclinorium), the successor basin formed in Early Silurian (Llandoverian) or Late Silurian (Pridolian) in northern Gaspe´ Peninsula and southern Que´bec, respectively. The post-Taconian basins are limited by two major orogenic events, the Middle-Late Ordovician Taconian and the Middle Devonian Acadian (Figure 3). A signiﬁcant Silurian deformation event is recognized along segments of the Appalachians and has been called the Salinian or Salinic Orogeny. However, Boucot (1962) introduced the term ‘‘Salinic’’ to designate a Late Silurian unconformity in northeastern USA. Over the years, confusion has arisen about the meaning of Salinic. The Salinian or Salinic Orogeny (Dunning et al., 1990; Cawood et al., 1994; van Staal and de Roo, 1995; Waldron et al., 1998; Tremblay and Castonguay, 2002; van Staal, 2005; Ettensohn, Chapter 4, this volume) is a Silurian event that resulted from the accretion of Ganderia along the outboard segment of the composite Laurentia margin (van Staal, 2005). In the distal part of the Quebec reentrant, its expression is locally subtle (Malo, 2001; Castonguay and Tremblay, 2003; Lavoie and Asselin, 2004; Tremblay and Pinet, 2005).

5.1. Newfoundland The post-Middle Ordovician successions in western Newfoundland form various depositional belts (Williams, 1995). The focus will be on the Clam Bank Belt because it is one of the few well-dated Middle Paleozoic successions of western Newfoundland, and its evolution can be tied into that of the adjacent Gaspe´ Belt (Figure 2). 5.1.1. The Clam Bank Belt This post-Middle Ordovician succession occurs in the Port au Port Peninsula and consists of limestone and siliciclastic of Late Ordovician (latest Darriwilian-Caradocian) to end Early Devonian (Emsian) age. Three distinct rock units are preserved: the Upper Ordovician Long Point Group, the Lower Devonian Clam Bank Formation and the upper Lower Devonian Red Island Road Formation (Figure 13). The Long Point Group (Riley, 1962; Bergstro¨m et al., 1974) represents the youngest Ordovician deposits in western Newfoundland. Until recently, the basal contact of the Long Point Group was not exposed in ﬁeld section and the Long Point was traditionally interpreted to unconformably overlie the Humber Arm Allochthon (Rodgers and Neale, 1963; Stevens, 1970; James and Cuffey, 1989; Stait and Barnes, 1991). Stockmal and Waldron (1990) and Waldron and Stockmal (1991), however, have argued, based on seismic information, that the Long Point Group is structurally at the top of a triangle zone and was thrusted easterly (Tea Cove Thrust) over its actual position (Waldron et al., 1993; Stockmal et al., 1995, 1998) (Figure 5). Recently, the contact between the two units was excavated and documented to be an erosional unconformity, later modiﬁed by folding (Batten and Dix, 2004). Waldron et al. (1998) interpreted the Long Point Group as belonging to a Late Taconian foreland basin on the basis of ongoing signiﬁcant subsidence that followed the Middle Ordovician ﬂysch sedimentation. The Long Point Group consists of three formations, the Lourdes, Winterhouse and Misty Point (Bergstro¨m et al., 1974; Quinn et al., 1999; Batten and Dix, 2004) (Figure 13). The upper Darriwilian-Caradocian Lourdes Formation consists of a lower thin assemblage of peritidal sandstone and limestone; the bulk of the unit however, is predominantly an open marine, nodular limestone with shale, thick calcarenite and calcirudite and small coral boundstones (James and Cuffey, 1989). The Caradocian Winterhouse Formation consists of limy sandstone, limestone conglomerate, siltstone and shale. Sedimentary facies and ichnofacies indicate a storm-dominated shelf (Quinn et al., 1999). The Caradocian Misty Point Formation (Quinn et al., 1999) consists of marginal-marine to terrestrial cross-bedded sandstones. Ross and Ross (1988) recognized a global eustatic sea-level lowstand at the Darriwilian-Caradocian boundary; this was followed by a complete eustatic T-R cycle in Caradocian time (see earlier section and Figure 13). The outer-shelf Caradocian Winterhouse Formation overlies the peritidal to shallow subtidal upper Darriwilian Lourdes Formation. The Long Point Group ends in the Caradocian Misty Point Formation which is dominated by marginal-marine to sub-aerial deposits, thus deﬁning a complete T-R cycle (Figure 13). Even if the Upper Ordovician Long Point Group is a Late Taconian foreland-basin ﬁll (Stockmal et al., 1995; Waldron et al., 1998; Quinn et al., 1999), sedimentation was likely primarily controlled by eustatic sea-level ﬂuctuations. The Clam Bank Formation (Rodgers, 1965; Bergstro¨m et al., 1974) unconformably overlies the Long Point Group (Quinn et al., 1999) (Figure 13). The lowermost Devonian Clam Bank Formation consists of crossbedded red sandstone and siltstone, with variegated shale and minor fossiliferous limy siltstone. The stratigraphic succession of the Clam Bank Formation indicates marginal-marine ﬂooding over the unconformity followed by rapid shallowing upward to terrestrial sedimentation (Morin, 1986; Burden et al., 2002). This T-R cycle matches that of the Lochkovian eustatic record of Dennison (1985) (Figure 13). The Red Island Road Formation was introduced by Williams et al. (1996) and elevated to formal status by Quinn et al. (2004). It consists of volcanic-rich (rhyolite) coarse conglomerate and sandstone for which an Emsian

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(late Early Devonian) age is proposed (Williams et al., 1996; Stockmal et al., 1998; Quinn et al., 2004) (Figure 13). The Emsian Red Island Road Formation is structurally deformed. It represents the last known sedimentary unit deposited in the Anticosti foreland basin.

5.1.2. The Salinic unconformity and orogeny in western Newfoundland A major unconformity is present between the Long Point Group and the Clam Bank Formation (Waldron et al., 1998) (Figure 13) with all Silurian strata missing. The unconformity is documented in offshore seismic imaging and the Clam Bank Formation unconformably overlies the Llandoverian succession of the Anticosti Group (Sanford and Grant, 1990). The strata missing render it difﬁcult to unequivocally correlate the unconformity with either the Late Silurian salinic unconformity (Boucot, 1962), which resulted from a global sea-level lowstand in Ludlovian-earliest Pridolian (Ross and Ross, 1996) or with the Silurian tectonic accretion of Ganderia (Salinic Orogeny, van Staal, 2005). The onset of sedimentation in earliest Devonian time (Clam Bank Formation) correlates with the post-unconformity eustatic-tectonic transgressive event recognized in the nearby Gaspe´ Peninsula (T2 event of Bourque et al., 1995, 2000, 2001; see further). Tectonic model in western Newfoundland suggests that Late Ordovician to earliest Silurian subsidence, deformation, burial and metamorphism followed Early Taconian thrusting (Dunning et al., 1990; Cawood, 1993; Cawood et al., 1994, 1995; Waldron et al., 1998; Stockmal et al., 1998; van Staal et al., 1998; van Staal, 2005). Tectonic subsidence was halted at the end of the Early Silurian followed by thermal and/or tectonic uplift, erosion and starved sedimentation (Waldron et al., 1998).

5.2. Continental Eastern Canada The end of Taconian ﬂysch sedimentation and/or emplacement of Ordovician oceanic units over continental slope allochthons in eastern North America mainland occurred diachronically from Middle Ordovician at or near promontories (St. Lawrence and New York) to Late Ordovician for the inner segment of the Que´bec Reentrant (Figure 13). The onset of sedimentation for various successor basins is also diachronous in a northwesterly trend. However, a signiﬁcant stratigraphic gap is reported along strike in the orogen and the sedimentary record in southwestern Que´bec is less complete compared to other post-Taconian basins. The post-Taconian successions for the continental segment of the Appalachian Orogen are collectively known as the Gaspe´ Belt (Bourque et al., 1995, 2000). The Middle Paleozoic (Late Ordovician to early Middle Devonian) Gaspe´ Belt is preserved in a number of tectonostratigraphic elements, including from north to south: the Connecticut Valley-Gaspe´ synclinorium, the Aroostook-Perce´ anticlinorium and the Chaleurs Bay synclinorium. The Gaspe´ Belt was folded and faulted with transpressive, dextral, strike-slip faults as the ﬁnal event (Malo and Be´land, 1989; Malo and Bourque, 1993; Kirkwood and Malo, 1993; Malo et al., 1995; Malo, 2001). Palinspastic reconstruction of the belt in eastern Que´bec has allowed reconstructing the regional paleogeography based on the restored geometry of the basin (Kirkwood, 1999; Bourque et al., 1995, 2000). From these reconstructions, the segment of the Gaspe´ Belt that forms southern Gaspe´ Peninsula was located above the current position of Cape Breton Island, therefore adjacent to the St. Lawrence Promontory (Figure 14). The

Figure 14 Palinspastic restoration of the Gaspe¤ depositional basin (blue-colored area) in eastern Que¤bec based on the work of Bourque et al. (1995, 2000, 2001) and Kirkwood (1999). The restoration eliminates signi¢cant dextral displacement along major Acadian transpressional faults as well as up to 50% of shortening accommodated by folds. Arrows indicate the general displacement along Acadian transpressional dextral faults.

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Figure 15 Simpli¢ed geological map of eastern Quebec and northern New Brunswick showing the distribution of the major group units of the Late Ordovician to Middle Devonian Gaspe¤ Belt. The major tectonostratigraphic domains are illustrated; the Connecticut Valley-Gaspe¤ synclinorium (CVGS), the Aroostook-Perce¤ Anticlinorium (APA) and the Chaleurs Bay Synclinorium (CBS). Geological map modi¢ed from Bourque et al. (1995).

summary of Bourque et al. (1995) presents paleogeographic and paleotectonic reconstructions of the Gaspe´ Belt and is still the reference work for correlation with adjacent Middle Paleozoic basins. On the northern side of the Gaspe´ Peninsula and in southern Que´bec (Figures 15 and 16) the Middle Paleozoic Gaspe´ Belt unconformably overlies Taconian Humber and Dunnage zones. In central and southern Gaspe´ Peninsula and in northern New Brunswick, the oldest sediments of the Gaspe´ Belt paraconformably overlie the Dunnage Zone rocks (Figure 15). However, locally, the Silurian section locally unconformably overlies outliers of the Humber and Dunnage zones (Maquereau and Mictaw groups in southern Gaspe´ and the Fournier and California Lake groups, northeastern New Brunswick). The post-Acadian Maritimes Basin (Lynch, 2001; Gibling et al., Chapter 6, this volume) has its depocenter in the Gulf of St. Lawrence and regionally, uppermost Devonian to Carboniferous sediments unconformably overlies the Gaspe´ Belt (Jutras et al., 1999, 2001). Within the three major tectonostratigraphic elements, three rock packages are proposed (Malo and Bourque, 1993; Bourque et al., 1995; Wilson et al., 2004; Lavoie and Asselin, 2004; Malo, 2004) (Figure 17). These assemblages are the results of three distinct regressive phases (R1–R3) separated by two transgressive events (T1 and T2) (Bourque et al., 1995, 2000; Figure 17).

5.2.1. Late Ordovician–Early Silurian (the R1 event) The Upper Ordovician (Caradocian/Ashgillian)–Lower Silurian (Wenlockian) package (Figure 17) occurs in all tectonostratigraphic elements. The stratigraphic record from northern New Brunswick to northern Gaspe´ Peninsula documents a northerly directed diachronous Late Ordovician to Early Silurian resumption of postTaconian sedimentation (Malo, 2004) (Figures 13 and 17). Upper Ordovician-Lower Silurian strata are unknown in southern Que´bec.

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Figure 16 Simpli¢ed geological map of southern Que¤bec with the distribution of post-Taconian units in the Connecticut Valley--Gaspe¤ synclinorium. The units in the green-shaded area (Lac Lambton (L), Ayers Cli¡ and Compton formations of the Saint-Francis Group and the Frontenac Formation) were displaced northwesterly over an unknown distance, along a major Acadian reverse fault (La Guadeloupe Fault). Middle Devonian intrusives are in red. Modi¢ed fromWilliams (1995).

For all localities other than northern Gaspe´ Peninsula (Figure 17), deep-marine clastics (Caradocian to Ashgillian Honorat Group and equivalent units) are overlain by below-wave base limestones (Ashgillian to Aeronian Matape´dia Group and equivalent units). The end of the ﬁrst regressive event is expressed by deep- to mid-shelf to nearshore clastics (Aeronian to Telychian Chaleurs Group) and the shallowing event ended in peritidal-dominated carbonate platform (Wenlockian Chaleurs Group). Stratigraphic sections on the northern edge of the Gaspe´ Belt overlie the post-Taconian unconformity; marine sedimentation only resumed in mid-Llandoverian (Figure 17). In northern Gaspe´ Peninsula, the base of the succession consists of outer-shelf clastic-limestone facies (Aeronian-Telychian Chaleurs Group) that passes upward to nearshore clastic and peritidal limestone (Telychian-Sheinwoodian Chaleurs Group). Tectonism played a key role in controlling the sea-level evolution after the end of Taconian foreland basin sedimentation. Ross and Ross (1988, 1996) proposed an overall regression for most of Ashgillian time followed by punctuated transgressive episodes in Llandoverian to slightly regressive conditions in Early Wenlockian (Sheinwoodian) (Figure 17). The curve of Johnson et al. (1998) generally follows the one of Ross and Ross (1996), although the Johnson et al. (1998) curve suggests transgressive conditions in Sheinwoodian. In Gaspe´, Bourque (2001) documented overall regressive conditions (the R1 event) for the ﬁrst stratigraphic package (Figure 17). The base-level rise documented by Bourque et al. (1995, 2000) and Bourque (2001) is coeval with the Early Silurian tectonic and/or thermal doming and uplift in western Newfoundland.

5.2.2. Early Silurian–Late Silurian (the T1-R2 events) The Lower (end-Wenlockian)–Upper (Pridolian) Silurian package forms part of the Chaleurs Group (Figure 17). The Salinic unconformity is present at the top of that stratigraphic interval (Bourque et al., 1995, 2000; Wilson, 2000, 2001, 2003a; Wilson et al., 2004; Lavoie and Asselin, 2004). A second-order sea-level rise is recorded in the deeper marine facies that overlies the Wenlockian carbonate platform (Figure 17) and culminates in the Ludlovian deeper marine clastics on the Gaspe´ Peninsula (Lavoie et al., 1992; Bourque, 2001). The rapid relative sea-level rise (T1 of Bourque et al., 1995; Bourque, 2001) is Late Wenlockian (Homerian) and is interpreted to have been partly controlled by extensional faulting (Lavoie et al., 1992; Lavoie and Morin, 2004; Lavoie and Chi, 2006). The following regressive succession is correlated with a second-order sea-level fall that started in Ludlovian and ended in a major Pridolian sea-level lowstand (R2 of Bourque et al., 1995; Bourque, 2001) (Figure 17). Thirdorder T-R cycles are preserved in carbonate facies developed in the late stages of the second-order shallowing event (West Point Formation reefal facies, Bourque et al., 2000; Bourque, 2001). The end of the regressive phase resulted in local sub-aerial exposure of the pre-Upper Silurian succession and deep erosion (Bourque et al., 2000); this

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Figure 17 General stratigraphic framework (at group level) for the post-Taconian units in the Que¤bec Reentrant. The detailed formation nomenclature (limits shown in thin dashed lines) can be found in Bourque (1975), Bourque and Lachambre (1980) and Bourque et al. (1993, 1995, 2000, 2001) for Gaspe¤ and Te¤miscouata, in Lavoie and Bourque (1992) and Lavoie and Asselin (2004) for southern Que¤bec and inWilson et al. (2004) and Lavoie and Asselin (2004) for New Brunswick. The major T-R events discussed in the text are shown; the Late Ordovician-Early Silurian R1 event; the Early Silurian-Late Silurian T1-R2 events and the latest Silurian-Middle Devonian T2-R3 events. The Silurian (R1-T1-R2 events) sea-level curve of Ross and Ross (1988, 1996) and the Devonian (T2-R3 events) one of Dennison (1985) are shown for comparative purposes. The red-colored areas indicate a time hiatus. Rhu., Rhuddanian; Aero., Aeronian; Tely.,Telychian; Shein., Sheinwoodian; Hom., Homerian; Lochk., Lochkovian; CBS, Chaleurs Bay synclinorium; APA, Aroostook-Perce¤ anticlinorium; CVGS, Connecticut Valley--Gaspe¤ synclinorium; NB, New Brunswick; LST, Limestones; SST, Sandstones; Re, reefal units discussed in text.

generated the Salinic unconformity at speciﬁc localities in the Gaspe´ Peninsula (Lachambre, 1987; Lavoie et al., 1992; Bourque et al., 2000; Bourque, 2001; Malo, 2001; Lavoie and Morin, 2004). As a whole, the T1-R2 succession matches relatively well the published second-order, eustatic, sea-level curves (Ross and Ross, 1996; Johnson et al., 1998) (Figure 17). However, the T1 event in the Gaspe´ Belt was ampliﬁed by active synsedimentary collapse of the depositional basin during a eustatic sea-level rise. The following R2 event was signiﬁcantly controlled by a major eustatic sea-level fall, the magnitude of which largely exceeded the still ongoing tectonic foundering of the depositional setting. 5.2.3. Latest Silurian–Middle Devonian (the T2-R3 events) In Gaspe´ Peninsula and New Brunswick, the Salinic unconformity commonly marks the base of the uppermost Silurian (Pridolian) to Middle Devonian (Eifelian) stratigraphic interval (Bourque et al., 1995, 2000; Wilson, 2000, 2001, 2003a; Wilson et al., 2004; Lavoie and Asselin, 2004). This rock assemblage is composed of the uppermost Silurian to lowermost Devonian section of the Chaleurs Group, followed by the Lower Devonian Upper Gaspe´ Limestones/Fortin/Dalhousie groups, and capped by the Lower Devonian to lowermost Middle Devonian Gaspe´ Sandstones Group (Figure 17). The preserved sedimentary record in the Gaspe´ Belt of southern Que´bec is no older than Late Silurian (Figure 17). A Late Silurian (Saint-Francis and correlative units; Figure 17) to Middle Devonian (Frontenac Formation; Figure 17) T-R assemblage is also recognized. The T2 event started in latest Silurian, with the development of major reef platforms above the Salinic unconformity (West Point, Laplante, Lac Aylmer, Sargent Bay formations; Lavoie, 1985; Bourque et al., 1986;

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Hughson and Stearn, 1989; Lavoie and Bourque, 1992; Bourque, 2001; Wilson et al., 2004; Figure 17). Locally, in southern Que´bec, a thick braided ﬂuvial conglomeratic succession overlies the basal unconformity (Lavoie, 1985; Lavoie and Bourque, 1992; Lavoie and Asselin, 2004). The West Point Formation in southern Gaspe´ built a thick (up to 600 m) reef margin in a short period of time (Pridolian) indicating combined signiﬁcant tectonic subsidence and eustatic sea-level rise (Bourque, 2001). In northern Gaspe´, the rapid sea-level rise resulted in the development of earliest Devonian isolated pinnacle reefs over the rapidly collapsing bioconstructed platform margin (Bourque, 2001). The platforms are overlain by earliest Devonian pro-deltaic to outer-shelf facies (uppermost Chaleurs and lower St. Francis Group; Bourque et al., 2000, 2001; Wilson, 2003b; Wilson et al., 2004; Lavoie and Asselin, 2004). These units record the increase in sea-level rise in the basin, which started in latest Pridolian (southern Gaspe´-northeastern New Brunswick) to earliest Lochkovian (northern Gaspe´) (T2 of Bourque et al., 1995) (Figure 17). The deepest marine conditions in the T2 event are recorded in the Lower Devonian (upper Lochkovian-Emsian) package consists of the mixed deep-marine limestone, clastic and volcanic facies (Upper Gaspe´ Limestones and equivalent units; Figure 17). The geographic distribution of these units indicates that a deep (below-wave base) marine carbonate ramp (Upper Gaspe´ Limestones) developed at the northeastern end of the Gaspe´ Belt (Lavoie, 1992a, 1992b). This carbonate ramp passes basinward (toward the south–southwest) into a deep, siliciclastic, marine environment with signiﬁcant intra-plate volcanic ﬂows and volcaniclastics (Fortin, Dalhousie and Saint-Francis groups) (Lavoie et al., 1991; Lavoie, 1992a, 1992b, 1995b; Bourque et al., 2000; Wilson, 2003b; Lavoie and Asselin, 2004). The overlying Lower (Emsian) to lower Middle (Eifelian) Devonian succession records a rapid tectonically controlled shoaling event. The Gaspe´ Sandstones Group (and equivalent units in northern New Brunswick; Cant and Walker, 1976; Rust, 1981; Desbiens, 1992; Bourque et al., 2000; Wilson et al., 2004; Figure 17) and the Frontenac Formation in southern Que´bec (Lavoie, 2004; Lavoie and Asselin, 2004) offer facies indicative of initial marginal-marine to ultimately delta plain and proximal braided-plain deposits. However, a wide area between Gaspe´ Peninsula and southern Quebec remained under deep-marine conditions. The T2-R3 events poorly match the eustatic curve, this indicates an overriding tectonic control on relative sea-level evolution. Dennison (1985) curve suggests relative third-order sea-level standstill to slightly regressive conditions at a time of the T2 event (Figure 17). Extensional tectonism occurred in Early Devonian time in eastern Gaspe´ (Lavoie, 1992a; Malo, 2001) and the rapid sea-level rise recorded at the base of the sequences could reﬂect such collapse. The sea-level rise was abruptly stopped in late Early Devonian; a rapid shoaling of facies is observed and resulted in the R3 event in the basin. This abrupt shallowing succession was deposited at a time of eustatic sea-level rise (Johnson et al., 1985; Dennison, 1985) (Figure 17) therefore, the observed regressive succession resulted from the building of the Acadian orogenic wedge (Malo, 2001). 5.2.4. The Salinic unconformity and disturbance in the Gaspe´ Belt In the Gaspe´ Belt, the presence of a Late Silurian erosional event has long been known in northern and southern Gaspe´ (Bourque et al., 1986; Lachambre, 1987; Lavoie, 1988; Bourque, 1990; Lavoie and Morin, 2004) and more recently in Northern New Brunswick (Wilson, 2001, 2002, 2003a; Wilson et al., 2004). The tectonostratigraphic work of Bourque et al. (1995, 2000, 2001) documented that this erosion resulted from sea-level lowstand (Figure 18). The oldest evidence of extensional collapse in the post-Taconian basin is found in the upper Llandoverian facies, which marks the inception of the Acadian foreland basin (Kirkwood et al., 2002, 2005). From late Early Silurian to Early Devonian, transtensive fault movements and early NW-folding occurred in the northern segment of the Gaspe´ Belt (Malo, 2001) and in northwestern New Brunswick (Carroll, 2003). These tectonic events were used to deﬁne the Salinic disturbance (Bourque, 1990; Bourque et al., 2000, 2001; Malo, 2001). Evidence from western Newfoundland (Waldron et al., 1998), central New Brunswick (van Staal and de Roo, 1995) and southern Gaspe´ (Kirkwood et al., 2002, 2005) indicate building of a tectonic wedge at or near the St. Lawrence Promontory in Early Silurian and the accretion of Ganderia on the Laurentia margin (van Staal, 2005). This event created a major Silurian foreland basin in the inner segment of the Que´bec Reentrant (Kirkwood et al., 2002, 2005). The tectono-sedimentary model is similar to what has been proposed by Bradley et al. (1998) for the adjacent and coeval Maine Appalachians. On Gaspe´ Peninsula, the extensional features ascribed to the Salinic disturbance are interpreted as the northern distal foreland of the tectonic wedge to the south (Figure 18). From recent seismic lines in the Gaspe´ Belt of eastern Quebec, the sub-surface geometry differs signiﬁcantly from the surface geology of large synclines and tight anticlines. A complex compressive structural style of reverse faults, blind thrusts, duplexes and triangle zone affects pre-Upper Silurian units (Morin and Laliberte´, 2002a, 2002b; Kirkwood et al., 2002, 2005; Figure 19). This geometry results from major compressive deformation of pre-existing strata (Lower to Upper Silurian) overlain by mildly deformed younger strata (Upper Silurian to

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Figure 18 Paleogeographic scenario in Late Ludlovian time in eastern Canada. Emerged areas and erosion of pre-Late Silurian units are recorded for western Newfoundland, for the palinspastically restored southern Gaspe¤ (and northern New Brunswick) Belt, and at the northern edge of the Gaspe¤ Belt. The cross-section (located on the paleogeographic map) shows that the Early Silurian (T1: Llandoverian) tectonic exhumation (through tectonic-thermal doming) in western Newfoundland resulted in a signi¢cant base-level rise in the southern Gaspe¤ Belt. The development of the tectonic wedge in Early Silurian resulted in the formation of a foreland-basin within the Que¤bec Reentrant. In Late Ludlovian--Pridolian time (T2), the global eustatic sea-level lowstand resulted in erosion (hatched areas) of topographic highs and generated the Salinic unconformity. The following Early Devonian eustatic rise (T3) was enhanced by the continuation of extensional faulting.

Figure 19 Seismic section based on composite seismic lines VB-4 (a, b and c) and 6 in western Gaspe¤ Peninsula. The Gaspe¤ Belt succession can be seismically divided into two major structural packages, a lower one that comprises Upper Ordovician to Upper Silurian units that is characterized by major compressive structures (blind faults, reverse thrust and triangle zone) and an upper one that comprises Devonian units that are characterized by fault-bounded tight anticlines and open synclines. Modi¢ed from Morin and Laliberte¤ (2002a, 2002b) and Kirkwood et al. (2005). The location of the seismic lines is shown by the green line on Figure 2.

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Middle Devonian). The geometry suggests a compression-formed foreland-propagating basin (Stockmal et al., 2003). In southern Que´bec, Castonguay et al. (1997, 2001), Tremblay and Castonguay (2002) and Tremblay and Pinet (2005) proposed based on metamorphic, geochronology, seismic and structural evidence, that some faults within and east of the Humber Zone recorded Late Silurian–Early Devonian extension prior to the compressive Middle Devonian event (Acadian Orogeny). This Late Silurian extensional event is ascribed to the Salinian Orogeny (Tremblay and Castonguay, 2002; Tremblay and Pinet, 2005). The following rapid marine inundation was accelerated by the overall transgressive event recorded along Laurentia at that time (T2 event). However, the exact age of inception for the post-Taconian basin in southern Que´bec is unknown. The youngest Taconian unit is the Caradocian Saint-Victor Formation (Magog Group, Dunnage Zone) (Cousineau, 1990; Tremblay et al., 1995). However, a Late Silurian (Pridolian) age is documented for carbonates locally a few meters above the lower unconformable contact (Figure 17).

5.3. Paleogeographic reconstruction of the post-Taconian basins The paleogeographic and paleotectonic reconstruction of the post-Taconian Gaspe´ Belt has been the subject of comprehensive synthesis by Bourque et al. (1995, 2000, 2001a), Malo (2001, 2004) and Kirkwood et al. (2002, 2005). These palinspastic reconstructions (Kirkwood, 1999) were mostly concerned with the eastern Que´bec regions. The major conclusion of this restoration is that the post-Taconian Gaspe´ Belt in eastern Que´bec and northern New Brunswick matched the shape of the Que´bec Reentrant and the St-Lawrence Promontory (Figure 14). Such palinspastic restoration of the Gaspe´ Belt in northern New Brunswick and southern Que´bec is not available; therefore, reconstructions have to be taken with that limitation. The recent studies of the SilurianDevonian units in southern Que´bec (Lavoie, 2004; Lavoie and Asselin, 2004) and northern New Brunswick (Wilson et al., 2004) allow the development of new preliminary regional paleogeographic maps for the Late Silurian (Pridolian) to Early Devonian (Emsian) interval in eastern Canada.

Figure 20 Paleogeographic map for the latest Pridolian-earliest Lochkovian in eastern Canada. An Upper Silurian reef tract (Bourque et al., 1986, 1995, 2000; Hughson and Stearn, 1989; Lavoie and Bourque, 1992) is recognized from southern Gaspe¤ to southern Que¤bec. The inset illustrates the palinspastically restored geometry of the basin upon which the paleogeographic map is based. The exact position of the Acadian-transported units in southern Que¤bec is largely speculative.

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5.3.1. Latest Pridolian/earliest Lochkovian Silurian tectonism signiﬁcantly controlled sedimentation patterns for the Gaspe´ Belt in Gaspe´ Peninsula (Bourque et al., 2001; Malo, 2001). However, the development of the Salinic unconformity occurred because of the signiﬁcant magnitude of the sea-level fall recorded in Ludlovian–earliest Pridolian time. The following sea-level rise (T2 of Bourque et al., 2001) was possibly ampliﬁed by the ongoing Acadian foreland-basin extensional collapse. Figure 20 is built on the Late Pridolian paleogeographic maps of Bourque et al. (1995, 2000, 2001). It is shown that a latest Silurian–earliest Devonian red-bed coastal plain stretched from Gaspe´ Peninsula (Plage Woodmans Member of the West Point Formation; Bourque, 2001) to western Newfoundland (Lower Devonian Clam Bank Formation; Burden et al., 2002). The Pridolian reef tract (upper reef complex of the West Point Formation; Bourque et al., 1995) extended from Gaspe´ Peninsula to northeastern New Brunswick (Wilson, 2003b; Wilson et al., 2004) to the Te´miscouata region (Lajoie et al., 1968; Dansereau and Bourque, 2001) and ﬁnally, to southern Que´bec (Hughson and Stearn, 1989; Lavoie and Bourque, 1992; Lavoie and Asselin, 2004) (Figure 20). Late Silurian reefal facies are recognized as south as northern New York (Ettensohn, Chapter 4, this volume). In southern Que´bec, mid- to outer-platform carbonates and siliciclastics surrounded the reefal units along the northwestern edge of the preserved depositional basin (Lavoie and Asselin, 2004). Southeasterly, it is interpreted that the mid- to outer-shelf carbonate-siliciclastic regime passed to deep outer-shelf/slope sedimentation in earliest Devonian time (Lavoie and Asselin, 2004).

5.3.2. Middle Lochkovian The northern segment of the Gaspe´ Belt in Gaspe´ Peninsula continued to collapse in the Early Devonian, resulting in the establishment of the pinnacle reefs of the West Point in northern Gaspe´ (Bourque, 2001). In Middle Lochkovian, the carbonate regime changed to a deep-shelf siliciclastic environment with prograding (basinward) ﬁne- to coarse-grained pro-delta to delta front facies (Bourque et al., 1995, 2001) (Figure 21).

Figure 21 Paleogeographic map for the mid-Lochkovian in eastern Canada. The Gaspe¤ Belt was surrounded by emerged areas feeding prograding (arrows) deltaic units (Bourque et al., 1995, 2000; Lavoie and Asselin, 2004; Lavoie, 2004). The inset illustrates the palinspastically restored geometry of the basin upon which the paleogeographic map is based. The exact position of the allochthonous units in southern Que¤bec is largely speculative.

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In southern Que´bec, the latest Silurian reefs and carbonate platform were overlain by pro-deltaic to outershelf units of the Compton Formation (Figure 21). The lateral facies variation indicates that a southeasterly progradation of a proximal siliciclastic wedge occurred in the early stage of the Early Lochkovian T2 event. Ultimately, with the transgression, a deeper outer-shelf regime was established over the entire area (Figure 21).

5.3.3. Pragian–Early Emsian On the northeastern Gaspe´ Peninsula, Silurian-Early Devonian extension reversed into a compressive regime near the end of Upper Gaspe´ Limestones deposition (Pragian-Early Emsian) (Malo, 2001). Signiﬁcant shortening of the Gaspe´ Belt in the Gaspe´ Peninsula was accommodated through folding and transpressive faulting (Malo, 2001) (Figure 22). The sedimentary evolution of the basin was regressive. In response to ongoing Acadian orogenesis, marginal-marine to ﬂuvial sediments (Gaspe´ Sandstones) were deposited along the northeastern margin of the Gaspe´ Belt in Gaspe´ Peninsula and in northern New Brunswick (Figure 22). In the southwestern portion of this basin, deeper marine sedimentation (Fortin Group) prevailed (Figure 22). In southern Que´bec, Pragian-aged sediments occur in carbonate-clastic deep-shelf facies in the north (Ayers Cliff Formation; Figure 22) and in northwesterly prograding proximal deltaic facies (Frontenac Formation; Figure 22). The progradation of the deltaic facies has been ascribed to the building of the Acadian wedge to the south (Figure 22).

Figure 22 Paleogeographic map for the Pragian-early Emsian in eastern Canada. The Gaspe¤ Belt in southern Que¤bec was the site of extensive shallow-marine to marginal-marine deltaic sedimentation (Lavoie, 2004; Lavoie and Asselin, 2004) prograding over outer-shelf facies to the north. The northeastern extension in Gaspe¤ Peninsula recorded the end of outer-shelf carbonate sedimentation (Lavoie, 1992a, 1992b; Bourque et al., 1995, 2000) with the initiation of progradation of Acadian alluvial fans (Rust, 1981). A seaway (Bourque et al., 1995, 2000) was still opened. The inset illustrates the palinspastically restored geometry of the basin upon which the paleogeographic map is based. The exact position of the allochthonous units in southern Que¤bec is largely speculative.

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6. The Sea Level Record in the Lower to Middle Paleozoic Appalachians in Eastern Canada: Eustasy vs. Tectonism The evolution of Lower to Middle Paleozoic Appalachian basins in eastern Canada responded to changing relative sea level. The deep-marine successions preserved in the Cambrian-Ordovician section of the Appalachians imperfectly recorded the ﬁne-scale sea-level ﬂuctuations. However, the nature of the shallowmarine sediments shed into the deeper marine settings indirectly recorded this part of the continental-margin history. The synthesis of our knowledge for both the shallow- and deep-marine settings produces a comprehensive sea-level history. The post-Taconian to syn-Acadian basins preserved in the Appalachians consist of an assemblage of shallow- to deep-marine units, and the Middle Paleozoic, relative, sea-level history can be more easily reconstructed. This section summarizes the evolution of the Cambrian to Devonian basins in eastern Canada, the controls on sedimentation patterns, either eustatic, tectonic or both.

6.1. The Early Cambrian–Late Ordovician Humber Appalachians The successions from western Newfoundland to southern Que´bec recorded the initial tectono-sedimentary events that shaped up the Appalachians (Figures 1 and 3). The rift-to-passive-margin development has been demonstrated to be fairly similar for both areas and primarily controlled by eustasy, but minor divergences in the evolution on that continental margin are related to local tectonic activity in the proximity of failed rifts at the Quebec Reentrant (Figures 3 and 11). The preserved slope records allow recognizing major eustatic sea-level events (end Early Cambrian global lowstand, Middle to Late Cambrian Grand Cycles and the slightly diachronic Tremadocian–Arenigian/mid-Arenigian lowstands) (Figure 11). These events characterize the Sauk Sequence of Sloss (1963) and are recorded in the Appalachians (see also Ettensohn, Chapter 4, this volume). At the continental margin of Laurentia, the end of the Sauk Sequence coincides with the inception of signiﬁcant oceanic seaﬂoor subduction and the migration of a tectonic forebulge (Jacobi, 1981; Knight et al., 1991; Ettensohn, Chapter 4, this volume). The Tippecanoe Sequence (Sloss, 1963) in eastern Canada was deposited in a tectonically active environment. The Taconian Appalachian Humber foreland basin followed a strongly diachronic along-strike evolution (Figures 3 and 12). The foreland shelf at the St. Lawrence Promontory was short-lived, being restricted to the Darriwilian with rapid ﬂysch encroachment (Figure 12). The foundering of the shelf at the St. Lawrence Promontory occurred at a time of global sea-level fall, and both shallow- and deep-marine facies patterns responded to a dominant tectonic signal (Figure 12). In the Que´bec Reentrant, foreland marine shelves were built while the Taconian foredeep ﬂysch migrated toward the shallower settings of the reentrant (Figure 12). Even with documented active tectonism (Lavoie, 1994; Lavoie et al., 2003b; Lemieux et al., 2003; Sharma et al., 2003), the eustatic signal is detected in the evolution of the shallow-marine shelves (Dix, 2003) until tectonic collapse and burial under ﬂysch in Late Caradocian to Ashgillian (Figure 12). The Taconian accretion of volcanic arcs occurred diachronically along the continental margin of Laurentia (see also Ettensohn, Chapter 4, this volume). The ﬁnal emplacement of oceanic units on the continental margin of Laurentia in western Newfoundland only occurred in Middle Devonian. There, the youngest ﬂysch deposited above the Laurentia shallow-foreland shelf is Darriwilian. This ﬂysch is unconformably overlain by endDarriwilian shallow-marine carbonates. The youngest ﬂysch in the Quebec Reentrant is upper Middle Ordovician to middle Upper Ordovician in southern Que´bec and the Gaspe´ Peninsula, respectively. The Taconian foreland-basin ended prior to the latest Ordovician glacio-eustatic lowstand that marks the limit of the Tippecanoe I/II sub-sequences.

6.2. The latest Ordovician to Middle Devonian Acadian basins Sedimentation in post-Taconian basins took place throughout most of the Taconian tectonostratigraphic zones. The post-Taconian sedimentation lasted until the Middle Devonian. The succession covers the Tippecanoe II sub-sequence and ended near the base of Sloss’s (1963) Kaskaskia Sequence (mid-Lower Devonian to end Mississippian). Diachronic Taconian ﬂysch sedimentation and local obduction of oceanic rocks, is followed by an along-strike diachronic resumption of sedimentation (Figure 13). At the St. Lawrence Promontory, the DarriwilianCaradocian Long Point Group overlies Taconian ﬂysch. The Long Point Group belongs to a Late Taconian foreland basin and records a T-R cycle that correlates with the Caradocian eustatic signal (Figure 13). In Caradocian however, active tectonism has been demonstrated to control relative sea level and foreland platform foundering in the adjacent Que´bec Reentrant (Figure 12).

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The evolution of the post-Taconian Gaspe´ Belt is best constrained in the northeastern segment of the Que´bec Reentrant (Bourque et al., 1995, 2000, 2001). The succession results from three second-order regressive phases (R1–R3) separated by two transgressive events (T1 and T2) (Figure 17). The R1 event (Caradocian–Early Silurian) resulted from base-level uplift in the successor basin (Figure 17); the rise is related to tectonic-thermal doming and Salinic Orogeny at the St. Lawrence Promontory in Early Silurian (Waldron et al., 1998; van Staal, 2005). The R1 event ended at a time of global sea-level rise (Figure 17). The following T1 event has been associated by Bourque et al. (2000) to a poorly deﬁned Homerian (Upper Wenlockian) sea-level rise (Figure 17). At that time, the northern segment of the Gaspe´ Belt recorded the ﬁrst pulses of a new extensional tectonic environment, which was created in response to the initial building of the Acadian orogenic wedge near the St. Lawrence Promontory (Kirkwood et al., 2002, 2005). The magnitude of the T1 event was enhanced by the extensional tectonic regime. The R2 event ended in Late Silurian with sub-aerial exposure of the Gaspe´ Belt (Figure 17). The Salinic unconformity (Boucot, 1962) resulted from the global Late Ludlovian–Early Pridolian eustatic lowstand in the Acadian foreland basin (Figure 18). The following T2 event led to the initiation of sedimentation in the southern Que´bec extension of the Gaspe´ Belt (Figure 17). The inception of the T2 event correlates with a sea-level rise in Early Devonian time (Figure 17). The magnitude of the sea-level rise resulted from combined eustatic and continued collapse of the Gaspe´ Belt (Lavoie, 1992a; Tremblay and Castonguay, 2002). The R3 event culminated in the ﬁnal closure of Iapetus oceanic basin along Laurentia. From a dominant extensional regime in Early Devonian, the basin inverted into a compressional high in Pragian-Emsian (Figure 21). The regressive conditions culminated in ﬂuvial sedimentation during Acadian transpressive faulting and folding.

7. Hydrocarbon Potential of the Appalachian Basins Hydrocarbon exploration from offshore Mesozoic units of Atlantic Canada is a well-established priority for the industry. Exploration in the frontier Paleozoic sedimentary basins is less mature and has had minimal success due in part to the limited sub-surface understanding of the architecture of these basins (Lavoie and Bourque, 2001). However, production of oil and gas is recorded in all tectonostratigraphic elements forming the ancient continental margin of Laurentia in Canada.

7.1. Lower Paleozoic belts — Humber Zone in Que´bec The Lower Paleozoic Humber Zone of the Quebec Appalachians did not receive signiﬁcant attention until a late 1960s exploration seismic survey by Shell Canada, using a foothill-style play concept. This led to the successful drilling of the 7 Bcf Saint-Flavien gas ﬁeld (Figure 4; Be´land and Morin, 2000; Bertrand et al., 2003a). The best-known source rock are Upper Ordovician black shales with TOC values reaching 3wt% and HI up to 154 (He´roux and Bertrand, 1991). Good source rocks are also present in Middle Ordovician black shale units with TOC and HI values up to 5.5wt% and 306, respectively (Comeau et al., 2004). Lower Ordovician shales have a fair potential (TOC up to 1.6wt%; Bertrand et al., 2003b). Surface maturity data indicate a northeasterly decrease from the US border (sterile; RoW3%) toward Quebec City (oil window–condensate; Ro ¼ 0.74–1.3%) (He´roux and Bertrand, 1991; Comeau et al., 2004). A northeasterly increase is noted from the Quebec City area toward the Gaspe´ Peninsula (oil window to sterile) (Chi et al., 2000; Comeau et al., 2004). Transported burial maturation is indicated by signiﬁcant maturity jumps from one tectonic slice to another, and at the transition between the St. Lawrence Platform and the Appalachian basin (He´roux and Bertrand, 1991). The main reservoir target consists of hydrothermally altered, fractured, intervals in tectonic slices of shallowmarine platform rocks of the St. Lawrence Platform (Bertrand et al., 2003a; Lavoie et al., 2005). Secondary reservoir targets are the thick Cambrian-Lower Ordovician coarse-grained submarine fan deposits where secondary dissolution porosity is locally abundant.

7.2. Lower Paleozoic belts — Humber Zone in western Newfoundland The ﬁrst report of oil in Newfoundland goes back to 1812 with notice of ﬂoating oil on Parsons Pond in western Newfoundland. Reinterpretation of some 1970s seismic was published in the early 1990s (Waldron and Stockmal,

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1991; Figure 5). The new interpretation documented the presence of an unknown Foothills-type structural style and a triangle zone at the edge of the deformation zone (Stockmal and Waldron, 1990; Stockmal et al., 1998, 2004). Extensive seismic and exploration drilling led to the discovery of the Garden Hill oil ﬁeld in 1995 (Figure 5; Cooper et al., 2001). The Cambrian–Ordovician deep-marine black shales have high TOC and HI (10.35%, 759) values. Extracts of these shales have good biomarker correlation with the oils (Fowler et al., 1995). The surface samples from western Newfoundland show a south to north increase in maturation (Williams et al., 1998). A depth-related maturation increase is documented in exploration holes. In most wells, deeper units (4,000+ m) are still in the oil window. The available burial-history scenarios indicate that the source rocks entered in the oil window during the Acadian Orogeny and a signiﬁcant migration event occurred after the development of the hydrothermal dolostone reservoir in Early Devonian (Lane, 1990). The main reservoir target consists of Early Ordovician intertidal to shallow subtidal facies. The best reservoirs occur where hydrothermal dolomites overprint earlier burial dolomites adjacent to, now inverted, extensional faults (Cooper et al., 2001). The lateral and vertical distribution of the reservoir is highly variable and relates to the proximity of the hydrothermal fairway. Secondary target reservoirs are the Lower Cambrian well-washed quartz arenites.

7.3. Lower Paleozoic belts — Gaspe´ Belt Exploration in Gaspe´ Peninsula started in the mid-19th century based on the presence of seeping oils near faults in Gaspe´ (Lavoie and Bourque, 2001). Regional seismic coverage in early 1980s led to ﬁrst geophysical-based drilling. Small reservoirs were recently put into production (Lavoie and Bourque, 2001; Lavoie et al., 2001b). The Lower Devonian shales have some fair to poor source rock potential with TOC values below 1.5 wt% (Bertrand and Malo, 2001; Roy et al., 2003). GC-MS and GC-IRMS studies of oil produced from Devonian reservoirs suggest that the potential source rock could be Ordovician black shales (Bertrand et al., 2003b; Roy et al., 2003). Maturation is highly variable in eastern Gaspe´ and northern New Brunswick, it ranges from locally immature to the dry gas zone; it also positively correlates with depth (Bertrand and Malo, 2001, 2004). The western sector of Gaspe´ Peninsula is characterized by higher thermal maturation values (Roy et al., 2003). Data suggest that both early and late hydrocarbon migrations occurred. Early migration from pre-Lower Silurian source rocks is recognized in Upper Ordovician to Lower Silurian units; migration occurred before the development of the regional Upper Silurian unconformity (Lavoie and Chi, 2002; Lavoie and Morin, 2004; Lavoie and Chi, 2006). The Upper Ordovician to Upper Silurian units of the Gaspe´ Belt have been involved in a foothill-style of deformation most likely related to the northward migration of the Devonian Acadian foreland basin (Figure 19). Gas reservoirs are found in Lower Devonian hydrothermally altered, fractured, carbonates (Lavoie et al., 2001b). High-quality oil is produced from coarse-grained units of the Lower Devonian sandstones. Good reservoir rocks are also documented in hydrothermal dolomites of the Lower and Upper Silurian carbonates (Lavoie and Chi, 2001; Lavoie and Morin, 2004; Lavoie, 2005; Lavoie and Chi, 2006).

REFERENCES Asselin, E., and Achab, A., 2004, Stratigraphic signiﬁcance of Lower Paleozoic chitinozoan assemblages from Eastern Canada. Canadian Journal of Earth Sciences, v. 41, pp. 489–505. Batten, K. L., and Dix, G. R., 2004, Unconformities, their signiﬁcance, and character within a Middle Ordovician carbonate platform: Lourdes Formation, western Newfoundland. Geological Association of Canada/Mineralogical Association of Canada Joint Annual Meeting, St. Catherines. Abstract on CD-ROM, 502 pp. Be´land, P., and Morin, C., 2000, Le gisement de gaz naturel de Saint-Flavien, Que´bec, Ministe`re des Ressources Naturelles du Que´bec, 19 pp. Bergstro¨m, S. M., Riva, J., and Kay, M., 1974, Signiﬁcance of conondonts, graptolites, and shelly faunas from the Ordovician of western and north-central Newfoundland. Canadian Journal of Earth Sciences, v. 11, pp. 1625–1660. Bernstein, L., 1992, A revised lithostratigraphy of the Lower-Middle Ordovician Beekmantown group, St. Lawrence Lowlands, Que´bec and Ontario. Canadian Journal of Earth Sciences, v. 29, pp. 2677–2694. Bernstein, L., James, N. P., and Lavoie, D., 1992, Cambro-Ordovician stratigraphy in the Que´bec Reentrant, Grosses-Roches-Les Me´chins area, Gaspe´sie, Que´bec, in Current research, Part E, Geological Survey of Canada, Paper 92-1E, pp. 381–392. Bertrand, R., and Malo, M., 2001, Source rock analysis, thermal maturation and hydrocarbon generation in the Siluro-Devonian rocks of the Gaspe´ Belt basin, Canada. Bulletin of Canadian Petroleum Geology, v. 49, pp. 238–261. Bertrand, R., and Malo, M., 2004, Maturation thermique, potentiel roche me`re des roches ordoviciennes a` de´voniennes du nord-ouest du Nouveau-Brunswick, Geological Survey of Canada, Open File 4886, 109 pp.

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CHAPTER 4

The Appalachian Foreland Basin in Eastern United States Frank R. Ettensohn

Contents 1. 2. 3. 4. 5. 6.

Introduction Appalachian Basin Elements and Limits Influence of Precambrian Events and Basement Paleogeographic/Paleoclimatic Framework Eustatic Framework Flexural Modeling of Foreland-Basin Sedimentation 6.1. Cycle origins 7. Generating the Appalachian Margin: Late Precambrian–Early Cambrian Rifting and Rift Fill (B765–B535 Ma) 8. The Appalachian Passive Margin 8.1. Latest Neoproterozoic–Early Ordovician pericratonic sedimentation (upper Sauk sequence, B570–472 Ma) 9. Two Orogenic Cycles and the Origin of the Appalachian Foreland Basin 9.1. Middle Ordovician–Early Devonian Caledonian orogenic cycle (Tippecanoe sequence, 472–411 Ma) 9.2. Early Devonian–Permian Variscan–Hercynian orogenic cycle (Kaskaskia and Absaroka sequences, 411–251 Ma) 10. Alleghanian Mountains, Post-Orogenic Collapse and Extension 11. Economic Resources and Potential 11.1. Energy resources 11.2. Coal-bed methane 11.3. Hydrocarbons 11.4. Carbon sequestration 11.5. Mineral resources 12. Discussion and Summary Acknowledgments References

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Abstract The Appalachian Basin is a composite, retroarc foreland basin that in many ways is the ‘‘type’’ foreland basin and the ‘‘type area’’ for the Wilson cycle. Our understanding of the basin, and others like it worldwide, is largely the legacy of a single observation by James Hall in 1857, an observation that also effectively established the framework for the later platetectonic paradigm, not to mention major framework developments in structure, tectonics, isostasy, flexural modeling, stratigraphy, sedimentation, paleoclimate and paleogeography. As preserved today, the basin is about 2,050 km long with an area of nearly 536,000 km2, extending from southern Quebec in Canada to northern Alabama in the U.S., and reflects the structural influence of earlier Grenvillian convergence and Rodinian dispersal, as well as the paleoclimatic, paleogeographic, eustatic and tectonic history of eastern Laurentia/Laurussia from latest Precambrian to Early Mesozoic time. During latest Precambrian to Early Ordovician time, the recently formed, southern to southeastern, Appalachian margin of Laurentia experienced mainly synrift and postrift, passive-margin sedimentation, largely controlled by local structure, regional climate and eustasy. However, by Cambrian time on some of the more distal, outboard parts of the Laurentian margin, the initial tectonic reorganization that would ultimately produce the Appalachian foreland basin had already begun. Major development of the Appalachian foreland basin began with the advent of the Taconian orogeny at about 472 Ma near the Early–Middle Ordovician transition and continued for nearly 200 Ma during four nearly continuous orogenies that reflect closure of the Iapetus and Rheic oceans and growth of Pangea. Tectonic dynamics controlled the extent and shape of the basin during various orogenies, and the resulting deformational loading is largely thought to have generated the Sedimentary Basins of the World, Volume 5 ISSN 1874-5997, DOI 10.1016/S1874-5997(08)00004-X

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accommodation space in which Appalachian sediments accumulated. Sediment thicknesses up to 13,700 m accumulated in 13, third-order (106–107 years), unconformity-bound, cycles that are clearly related to Appalachian tectophases, distinct phases of tectonism controlled by sequential convergence with continental promontories during orogeny. Tectophase cycles during the Taconian and Salinic orogenies and during the succeeding Acadian and Alleghanian orogenies form the larger, second-order (107–108 years), Caledonian and Variscan–Hercynian orogenic supercycles, which generally reflect closure of the Iapetus and Rheic oceans, respectively. These supercycles are separated by the brief, Siluro-Devonian, Helderberg interval, which in the foreland basin is represented by a thin, widespread, shallow-water, clastic and carbonate succession with poorly developed tectophase cycles. In contrast to the relative tectonic quiescence apparent in the foreland basin, evidence from more outboard parts of the orogen indicates that the Helderberg interval appears to represent a transitional period of uplift, magmatism and successor-basin formation during Taconian–Salinic orogen collapse and change to collision-related, strike-slip and transpressional regimes in succeeding orogenies. The first 11 cycles in the foreland basin mainly reflect subduction-type orogenies and typically consist of basal, dark, marine shales succeeded in ascending order by flysch-like and molasse-like units, all of which track the progress of orogeny in time and space. The last two Alleghanian cycles, in contrast, reflect collision-type orogeny and are largely composed of clasticdominated, terrestrial or marginal-marine sediments with a strong eustatic overprint related to Gondwanan glaciation. Although Alleghanian tectonism probably continued through Late Permian time, no foreland-basin sediments younger than Early Permian age are preserved. By Late Triassic time, thermo-tectonic thickening and uplift in the Alleghanian orogen caused orogen collapse and extension, ending the Iapetan or Appalachian Wilson cycle and initiating Pangean dispersal and the current Atlantic Wilson cycle. The importance of the Appalachian Basin lies not only in its ‘‘type’’ status as a basis for our understanding of geomorphological, structural, stratigraphic and sedimentary parts of the plate-tectonic paradigm, but also in the fact that it contains the relatively well-preserved, 545 Ma, stratigraphic and sedimentary record of one complete Wilson cycle and parts of others. The larger foreland-basin/orogen area clearly shows the orogen collapse and extension phase of the previous Laurentian or Grenvillian Wilson cycle during dispersal of the Rodinia supercontinent, as well as late-synrift, passive-margin, active-margin and orogen collapse phases of the Iapetan or Appalachian cycle during accretion and dispersal of the Pangean supercontinent. The foreland-basin area itself shows evidence for all the phases except orogen collapse. Nonetheless, what is particularly apparent throughout the basin’s entire history is the fact that the zigzag shape of the old Iapetan margin and the basement structural framework remaining at the end of the previous Laurentian cycle, combined with a series of probably global tectonic events, essentially controlled development and infill of the Appalachian foreland basin. This is apparent in the timing and distribution of the 13 sedimentary cycles that largely comprise its sedimentary infill. Even so, every cycle in the basin differs, reflecting the indelible overprint of changing climatic, geographic and eustatic regimes.

1. Introduction A single observation can fundamentally change the way we construct large parts of a discipline, and so it was with James Hall’s 1857 observation that the Paleozoic rocks in a linear belt along the eastern coast of the United States were many times thicker than rocks of the same age in the Upper Mississippi Valley. Although others had made similar observations in other places (e.g., E´lie de Beaumont, 1829), and there can be little doubt that Hall was aware of these observations (S- eng+or, 2003), Hall (1857, 1859, 1883) was first to recognize the ‘‘basinal’’ aspect of the Appalachian Mountains area, and with this recognition, provided the impetus for understanding the genetic relationship between such borderland belts of great sediment thickness and their subsequent deformation into mountain belts, an idea that established the groundwork for the later plate-tectonic paradigm. Although Hall did not call the mountainous area a basin, his observation caused a groundswell of work that would lead to these types of conclusions by others. In fact, it remained for Dana (1873) to name these broad, linear, sediment-filled basins geosynclinals (later called geosynclines); for Williams (1897) to name the Appalachian area; and for Dana (1856, 1867) and Schuchert (1923) to provide possible explanations and historical context. Foreign workers like Haug (1900) and Stille (1936) further refined the ideas and made them international. The Appalachian area also became fertile ground for work on structural and tectonic concepts like fold belts (Rogers and Rogers, 1843), isostasy (Dutton, 1889), deformational mechanisms (Willis, 1893), thrusts (Rich, 1934) and ‘‘thin-skinned’’ deformation (Rodgers, 1949, 1963; Gwinn, 1964). As a result, the Appalachian Basin became the ‘‘type’’ geosyncline and the Valley and Ridge province became typical of ‘‘Appalachian-type’’ structure, serving as models for mountain borderlands across the world (Schuchert, 1923). Further discussion on the role of the Appalachian Basin in the development of these ideas and others can be found in work by Rodgers (1970), Mayo (1985), Rast (1989), Bally (1989), S- eng+or (2003), Friedman et al. (2003) and Sanders and Friedman (2003). Similarly, the Appalachian Basin became important in understanding how the interplay of sea-level, paleogeography and tectonics influenced basin-wide stratigraphy, sedimentation and facies distribution, and important developments

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in these fields from the Appalachian Basin include studies by Schuchert (1955), Sloss et al. (1960), Colton (1970), Dennison and Wheeler (1975) and Dennison (1986). By tectonic setting and origin, the Appalachian Basin is a multistage or composite, retroarc foreland basin (e.g., Dickinson, 1974; Miall, 1995) that largely formed in response to tectonic loading during four, nearly continuous orogenies on the eastern margin of Laurentia/Laurussia during approximately 220 Ma from the Early–Middle Ordovician transition through Permian time. In many ways, the basin is the ‘‘type area’’ of the Wilson cycle (Wilson, 1966; Dewey and Burke, 1974; see also Miall, Chapter 17) and reflects foreland-basin sedimentation during one Wilson cycle, or the opening and closing of the Atlantic-type, Iapetus and Rheic oceans (Wilson, 1966; Harland and Gayer, 1972; Dewey and Burke, 1974); it is also an example of craton-margin sedimentation during middle to later parts of a supercontinent cycle (Worsley et al., 1984, 1985, 1986; Nance et al., 1986) from the breakup of Rodinia to the formation of Pangea. The basin mostly developed on a Late Precambrian (Neoproterozoic)–Cambrian, extensional, ramp margin generated by the breakout of Laurentia from the supercontinent Rodinia, beginning about 750 Ma (Hoffman, 1991; Torsvik et al., 1996; Dalziel et al., 2000; Torsvik, 2003), to form the adjacent Iapetus Ocean during a period of time ranging from approximately 750–535 Ma. The ramp margin was largely formed on Grenvillian crust (1.0–1.2 Ga) with a veneer of Cambrian and Lower Ordovician clastic and carbonate sediments. Development of the Appalachian foreland basin began with the advent of the first Paleozoic orogeny near the Early–Middle Ordovician transition (B472 Ma) and continued for nearly 200 Ma as the large interior Appalachian orogen (Nance and Murphy, 1994) grew through closure of the Iapetus and Rheic oceans; it was tectonic loading (e.g., Beaumont, 1981; Quinlan and Beaumont, 1984) of the crust during the incremental growth of this interior orogen that largely generated the accommodation space occupied by the well-known Appalachian sedimentary record. How that space was infilled with sediments, moreover, was controlled by several factors, including nature of the Precambrian basement, paleogeographic/paleoclimatic framework, eustatic changes and the loading/relaxational (flexural) history of adjacent orogens, and each of these topics will be briefly reviewed below. In developing this framework and history, the recently revised geologic time scale of Gradstein and Ogg (2004) and Gradstein et al. (2004a, 2004b) will be used.

2. Appalachian Basin Elements and Limits If we assume that a basin is an area ‘‘in which sediments can accumulate to considerable thickness and be preserved for long geological time periods (Einsele, 2000),’’ clearly, the Appalachian area, as it stretches from Alabama to Newfoundland, contains several basins. In this chapter, we will mainly examine the Appalachian foreland basin, an oblong sedimentary basin extending from southeastern Ontario and southern Quebec to northeastern Alabama (Figure 1). As Colton (1970) defined the basin, it is bound on the west by the Cincinnati, Findlay, and Algonquin arches, and on the east by metasedimentary, metavolcanic, and intrusive Precambrian and Paleozoic rocks of the Adirondack dome, Blue Ridge and New England Uplands (Figure 1). Its current northwestern boundary in southeastern Ontario and southern Quebec is defined by the updip erosional limit of Paleozoic sediments along the Laurentian and Frontenac arches of the Canadian Shield (Figure 1; see Lavoie, Chapter 3), whereas its southern boundary is transitional into the Black Warrior Basin. Clearly, the eastern boundary is somewhat arbitrary, as seismic studies (Clark et al., 1978; Teglund, 1978; Cook et al., 1979, 1981; Milici et al., 1979; Harris et al., 1981) show that relatively undeformed, Paleozoic, basin strata extend below the Blue Ridge, and perhaps under the Piedmont. This definition also excludes Precambrian sedimentary rocks and metamorphosed Paleozoic rocks within or east of the Blue Ridge that some workers might include within the basin or would suggest are closely related to it. Where appropriate, implications from rocks beyond the foreland basin, especially those in New England, will also be included. Thus defined, the presently preserved Appalachian foreland basin is largely a product of the Alleghanian orogeny and is approximately 2,050 km (1,270 mi) long, about 530 km (330 mi) wide at it broadest point, and includes an area of about 536,000 km2 (206,900 mi2) (Colton, 1970) (Figure 1). The basin contains mostly Paleozoic strata, which are preserved to thicknesses of 600– 900 m on its western flank and to more than 13,700 m on its eastern flank in the thickest area of central Pennsylvania (Patchen et al., 1985a, 1985b; de Witt and Milici, 1989). The sedimentary volume is estimated to be about 2,300,000 km3 (500,000 mi3) (Burns and Claus, 1985). Preserved parts of the basin extend into southern Quebec in southeastern Canada on the St. Lawrence Platform, and at one time the basin extended farther to the north and northeast in Canada (Poole et al., 1968; Russell, 1984; Sanford, 1993b; Faill, 1997b; Lavoie, Chapter 3) (Figure 1). Isopachous maps of Paleozoic units, as well as Paleozoic inliers in the Adirondack Mountains (Fisher et al., 1970; Rickard, 1973), indicate that parts of the basin must have extended above the inboard, Precambrian, Adirondack massif (Figure 1), which experienced uplift and unroofing as early as Cretaceous time (Roden-Tice

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Figure 1 Location of Appalachian foreland basin and nearby structural features in eastern United States and southeastern Canada, where dotted lines represent modern political boundaries. Preserved parts of the Appalachian foreland basin (A) are shown in coarse stipple and other possible parts of the basin to the northeast and northwest are outlined in dark dashes.The BlackWarrior Basin (B) and the Illinois (I), Michigan (M), and Moose River (MR) intracratonic basins are shown in light stipple.The basins are separated from each other by cratonic arches (la, Laurentian arch; fr, Frontenac arch; al, Algonquin arch; f, Findlay arch; c, Cincinnati arch).The large dark triangles show the directions and locations of yoking between the Appalachian and adjacent intracratonic basins. Major Iapetan rifts (r, Rome trough; k, Knoxville graben; b, Birmingham graben), de¢ning the former keel of the basin, are shown in a sand-and-gravel pattern bound by tick-marked dashes. Reactivation of some former rifts and cratonic arches at times served to subdivide the larger Appalachian basin into sub-basins. Blackened areas east of the Appalachian Basin represent a chain of Grenvillian, Mesoproterozoic basement inliers or massifs, called the Blue Ridge in the south and central Appalachians, which were parts of the old continental margin that were transported westward during various Appalachian orogenies; the Adirondak massif (a) is an inlier of similar age that was not involved in Appalachian orogenies.The Piedmont or Piedmont Plateau is the area of largely metamorphic and igneous rocks between the Blue Ridge and the Coastal Plain overlap. Other major structural features include the Grenville Front (gf), southern gravity anomaly (ga), the Allegheny Front (af), the 38th Parallel Lineament (t), the Ontario Embayment (o), the Quebec Basin (q), the Narragansett Basin (n), and a series of Carboniferous--Mesozoic strike slip faults (nk, N40-Kelvin fault zone; cc, Cobequid-Chedabucto fault zone) along which smaller sedimentary basins developed. ME, Maine; Qu, Quebec; ON, Ontario; RI, Rhode Island; cpo, Coastal Plain overlap; and AL, Alabama.

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et al., 2000). Similarly, it is likely that parts of the basin once extended farther to the east, but are hidden below adjacent allochthonous crystalline rocks or lost due to tectonism, uplift and erosion. Isopachous maps and modeling also show that at times western parts of the Appalachian foreland basin and the present-day Michigan Basin area merged to form a single large foreland basin (Beaumont et al., 1988; Coakley and Gurnis, 1995; Ettensohn and Brett, 2002). Hence, the Appalachian Basin is not just the feature that we see today; during most of Paleozoic time, it was probably a nearly continuous, elongate basin, sometimes divided into sub-basins, that extended from Alabama to Quebec, and possibly to Newfoundland (Poole et al., 1968; Sanford, 1993b; Faill, 1997b), and at times as far west as Michigan. At other times, it is clear that the Appalachian Basin had yoked with the adjacent intracratonic Moose River, Michigan and Illinois basins via intervening structural lows (Figure 1), such that ‘‘Appalachian influence’’ extended far beyond the basin boundaries (Quinlan and Beaumont, 1984; Ettensohn and Geller, 1987; Beaumont et al., 1988; Ettensohn et al., 1989; Sanford, 1993b; Coakley and Gurnis, 1995; Ettensohn and Brett, 2002). Its extent and shape have varied through time depending upon tectonic dynamics; what we see of the basin today is a composite that mostly reflects Alleghanian orogeny and postAlleghanian erosion. Perhaps the most salient reference feature for the Appalachian area is the Blue Ridge and related, along-strike, basement massifs to the north (Figure 1). They are largely formed of allochthonous, Proterozoic (Grenvillian) crystalline rocks that were torn away from the Late Proterozoic–earliest Paleozoic, Laurentian continental margin and transported as much as 400 km westward onto the craton aboard thrust sheets; these rocks are Laurentiancycle basement precursors to Paleozoic, Iapetan or Appalachian, Wilson-cycle tectonism (Hatcher, 1987, 1999). Rodgers (1968) suggested that the eastern edge of external massifs like the Blue Ridge corresponded approximately to the easternmost position of the continental margin, but the former location of that nowconcealed margin can only be inferred from a prominent gravity gradient (Figure 1, ga) in the southern Appalachians (Hatcher and Zietz, 1980; Bartholomew and Lewis, 1992). Moreover, the western edge of the same along-strike belt of crystalline massifs (Figure 1) approximates the eastern edge of the unmetamorphosed foreland and of the preserved Appalachian Basin. Most rocks to the east of the Blue Ridge and the related Precambrian massifs are considered to be part of the metamorphic core or internides, the internal part of orogens that was subject to intense ductile deformation and plutonism. Physiographically, the internides largely correspond to the Piedmont Plateau province. Those parts of the Appalachian area that were subject to marginal deformation in the form of folding and thrusting are included in the Appalachian externides, which are subdivided into crystalline and sedimentary parts (see Faill, 1998). The crystalline externides include the Blue Ridge and related Precambrian massifs, and perhaps parts of the Piedmont; because of uncertainty about what underlies the Piedmont, however, it is included here within the internides. The sedimentary externides largely include the Appalachian Plateau, Valley and Ridge, and St. Lawrence Lowland (Ottawa-Quebec Lowland) provinces (Figure 2). The easternmost sedimentary externides include a strongly deformed fold-thrust belt of mostly unmetamorphosed Paleozoic sedimentary rocks, which largely coincides with the St. Lawrence Lowland and Valley and Ridge physiographic provinces (Lavoie, Chapter 3; Figure 2). Much of the folding is related to blind thrusts at depth, and Figure 2 shows schematically the hypothetical arrangement of major Alleghanian de´collement zones. Remaining parts of the sedimentary externides are included in the adjacent Appalachian Plateau physiographic province, which is underlain by the broad Allegheny synclinorium. The western boundary of the plateau is defined by the west-facing Highland Rim and/or Pottsville escarpments, whereas the eastern boundary is approximated by an east-facing escarpment called the Allegheny Front (Figures 1 and 2). The Appalachian Plateau is a highly dissected area characterized by flatlying or very gently deformed sedimentary rocks. Remaining parts of the Appalachian Basin as far west and northwest as the defining arches are included in the Interior Lowland and Great Lake Lowland physiographic provinces (see Fenneman, 1938; Lobeck, 1941; Rodgers, 1970; Sanford, 1993a).

3. Influence of Precambrian Events and Basement Geophysical data have demonstrated that the Appalachian Basin and related orogenic rocks to the east are parts of the larger and once mostly continuous Appalachian–Caledonian orogen on both sides of the Atlantic that in part developed on Late Precambrian Grenvillian crust (e.g., Haworth et al., 1988; Powell et al., 1988; Rankin et al., 1989, 1993). The Grenvillian crust of eastern Laurentia, which largely underlies the basin, has isotopic ages that fall between 1.35 and 0.95 Ga and is the product of diachronous, largely Mesoproterozoic metamorphism and deformation from Labrador to the Llano Uplift of east Texas (Rankin et al., 1993; Mosher, 1998); it reflects a new segment of crust added to proto-Laurentia in several orogenic and accretion events during culmination of the earlier Laurentian, Wilson cycle during which the supercontinent Rodinia formed (e.g., McLelland et al., 1996; Hatcher, 1999, 2005). This crust has been encountered in a few test wells throughout the basin, and is

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Figure 2 Schematic west-east cross-section from central Kentucky to central Virginia through the Appalachian sedimentary and crystalline externides, showing the hypothetical structural and stratigraphic arrangement of de¤collement zones formed by the Alleghany orogeny. The section also shows the Appalachian Basin and parts of the basin present below the adjacent Blue Ridge and other Precambrian massifs. The broad basin-like structure below the Appalachian Plateau is the Allegheny synclinorium. Symbols: thick numbered lines, de¤collement stratum; thinner unnumbered lines, conformable system boundaries; dashed stipple, Precambrian crystalline basement; undulatory line, unconformity atop crystalline basement; dotted stipple, Precambrian and Cambrian basal sandstones. De¤collement horizons: 1, Cambrian Rome Fm; 2, Cambrian Conasauga Shale; 3, Ordovician Athens-Sevier-Rockmart shales; 4, Ordovician bentonites and Moccasin Fm; 5, Ordovician Martinsburg Sh; 6, Silurian Salina evaporites; 7, Devonian Mandata Sh; 8, Devonian Tioga bentonite; 9, Devonian dark shales; 10, Mississippian Maccrady evaporites; 11, Mississippi Floyd Sh (Alabama); 12, Mississippian Mauch Chunk-Pennington fms; 13, Pennsylvanian shales. Adapted from Dennison (1984) and Lobeck (1941).

becoming much better known through geochemical, geochronological, structural, gravity, magnetic, geophysical and historical analyses of the basin basement, Adirondak Massif and various Appalachian outlier and inlier massifs to the south; some of the most recent work is presented in Rankin et al. (1993) and in a volume edited by Tollo et al. (2004a). One prominent magnetic gradient in the Grenville crust, the New York–Alabama lineament (Figure 3), extends nearly 1,700 km through the center of the Appalachian basin and has been interpreted to represent a crustal break or shear zone that may coincide with an intra-Grenville suture (King and Zietz, 1978; Thomas, 2006); it is coincident with a trend of persistent subsidence in the Appalachian Basin called the Appalachian Basin keel line (Dennison, 1982, 1984). In fact, the eastern margin of the current Appalachian Basin is defined by a series of outlier anticlinoria or massifs (Figure 1), which represent blocks of Grenville crust rifted away from the new continent during Rodinian breakup and later reaccreted, or parts of the then new Laurentian crustal margin and its cover that were thrust 100–400 km cratonward during later Paleozoic orogenies (e.g., Cook et al., 1979; Tollo et al., 2004b; Hatcher et al., 2004; Hatcher, 2005). Those massifs like the Blue Ridge and other related, along-strike, basement massifs, which occur at the edge of the thrust belt and separate undeformed sedimentary rocks to the west from metamorphic rocks to the east, have been called external massifs; they may still retain Grenvillian structural style and be covered with Neoproterozoic and/or Cambrian clastic sediments or metasediments (e.g., Hatcher, 1983; Hatcher et al., 1989, 2004; Rankin et al., 1989; Rast, 1989). Other Precambrian massifs exposed as dome-like structures farther to the east within the metamorphic core, or internides, have been called internal massifs; they commonly exhibit a sedimentary cover that is metamorphosed and intensely infolded with basement rocks. All of these basement massif rocks were apparently once part of the Grenvillian crustal segment added to an older continental nucleus. The preserved western margin of this segment, called the Grenville Front (Figure 1, gf ) in part coincides with the western margin of the Appalachian Basin, and reactivation of structures comprising it at times probably influenced the extent of the foreland basin. Uncertainty still exists about the nature of the Grenville event and the resulting supercontinent (e.g., Piper, 2000), but consensus is developing that it was part of a global orogenic event with several phases of orogeny involving the collision of various Gondwanan cratons with the eastern margin of proto-Laurentia, 1.3–0.9 Ga ago during formation of the supercontinent of Rodinia (Hoffman, 1991; Torsvik et al., 1996; Dalziel et al., 2000; Torsvik, 2003; Tollo et al., 2004a). Most of the collision with the proto-Laurentian margin probably occurred at low latitudes with much precipitation so that the Grenville orogen was rapidly unroofed, leading to rapid uplift, deep erosion and an overfilled foreland basin of low preservation potential (Miall, 1995). Although some

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Figure 3 Appalachian foreland basin (A) relative to promontories and intervening reentrants on the Late Precambrian--Early Paleozoic, Laurentian continental margin formed by Iapetan rifts (thin, notheast-oriented lines) and transform faults (thicker SE-oriented dashed lines) during the breakout of Laurentia from Rodinia. Coarse stippled blocks outboard of main rifted margin are possible stranded fragments or ribbon microcontinents of Grenvillian crust that were re-accreted in subsequent orogenies; most are now preserved as internal massifs (see Figure 1). The Humber seaway was a narrow body of water between the Dashwoods ribbon microcontinent and the Humber margin of Laurentia. Promontory location apparently controlled tectophase development during orogenies, areas of greatest subsidence and thickest sediment accumulation in the foreland basin (dark stipple), and the areas of greatest deformation and metamorphism, as in the anthracite basins (An). NYAL, New York-Alabama lineament. State and province abbreviations: AL, Alabama; GA, Georgia; SC, South Carolina; NC, North Carolina; TN, Tennessee; KY, Kentucky; VA,Virginia; MD, Maryland; WV,West Virginia; OH, Ohio; PA, Pennsylvania; NJ, New Jersey; NY, NewYork; CT, Connecticut; MA, Massachusetts; RI, Rhode Island; VT,Vermont; ME, Maine; ON, Ontario; QU, Quebec; NB, New Brunswick; NS, Nova Scotia; and NF, Newfoundland. Partially adapted from Thomas (1993, 2006) and Hatcher et al. (1989).

Grenville foreland-basin sediments may be preserved in the subsurface of the eastern Midwest (Drahovzal et al., 1992; Santos et al., 2002), the Grenville foreland basin and most of the orogen were subsequently razed such that parts of the Late Neoproterozoic–Early Cambrian Appalachian area must have been a relatively featureless, gently sloping plain that facilitated rapid transgression into the area by Sauk seas. The exact organization of Rodinia is still debated (e.g., Hartz and Torsvik, 2002), but the continent persisted for about 300 Ma, and most workers believe that future Laurentia occupied a keystone position near its center, and that fragmentation and Laurentian breakout had begun by 750–600 Ma and continued until at least 535 Ma (Condie, 1997; Cawood et al., 2001). This breakout, however, had several consequences for the resulting Appalachian margin that would influence later development of the Appalachian Basin. First, the breakout resulted in a curvilinear or zigzagging, scarp-like continental margin (Bird and Dewey, 1970) that developed an irregular pattern of continental promontories and embayments (convex and concave, respectively, toward the ocean)

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(Figure 3), because it developed along valleys between triple junctions (Rankin, 1976), or because of separation into rift-transform segments (Thomas, 1977, 1993, 2006). Not only did this pattern control location of structural recesses and salients (concave and convex respectively toward the craton) in the orogen (Thomas, 1983), and indirectly, the shape of successive foreland basins, but the projecting promontories also served to localize tectonism on and near them because of greater shortening and more intense deformation (Dewey and Burke, 1974; Dewey and Kidd, 1974). The greater intensity of deformational loading at these points also insured greater flexural subsidence in adjacent parts of the foreland basin so that the three major depocenters or sub-basins of the Appalachian Basin (Lafferty, 1941; Hatcher et al., 1989) are located just behind or adjacent to each promontory (Figure 3). Where preserved, thick sedimentary successions may have also been deposited within the continental reentrants on converging lithosphere in peripheral, forearc, successor or pro-foreland basins, but these successions were commonly destroyed or intensely metamorphosed by repeated tectonic processes on the internal side of successive orogenies. For example, more than 2.5 km of such sediments, the thickest in eastern Canada, are preserved in the Quebec reentrant (Figure 3; see Poole et al., 1968; Lavoie, Chapter 3). In addition, Ettensohn (1985a, 1985b, 1991, 1994) has suggested that the natural division of Appalachian orogenies into tectophases (e.g., Johnson, 1971) (Figure 4) was mediated by diachronous convergence with successive promontories, which controlled third-order, sedimentary cyclicity, major unconformity distribution and the occurrence of clastic wedges in the Appalachian Basin (Ettensohn, 1994, 2004, 2005). The prominent northwest-southeast-trending transforms that partially delineated the margin (Figure 3) were reactivated successively and segmented the orogen and basin into blocks that responded differently to tectonic events (e.g., Rast and Skehan, 1999); these cross-strike faults also affected basin sedimentation (e.g., Castle, 2001), later basin mineralization (e.g., Heyl, 1972), hydrocarbon accumulation (Shumaker, 1996) and even final orogen collapse (Manspeizer, 1994). During Neoproterozoic to Early Cambrian time, aulacogens and ensialic shelf basins apparently developed along the margin, whereas intracratonic, Newark-like, rift basins paralleled the continental margin and developed some distance inland as parts of the Rome trough rift system (DeWindt, 1975; Shumaker, 1986a, 1996; Schwab et al., 1988; Rankin et al., 1989; Shumaker and Wilson, 1996; Gao et al., 2000; Gates and Volkert, 2004) (Figure 1). These basins were largely infilled with coarse terrestrial clastics, diamictites and volcanics, but in some basins, marine continental-slope, as well as inland-sea, continental-shelf and rise deposits, appear to have developed (DeWindt, 1975; Schwab et al., 1988; Rast and Kohles, 1986). The most important aspect about these basins, however, is the fact that the structures bounding them were periodically reactivated during subsequent orogenies and commonly controlled local patterns of sedimentation and later detached structural trends in the Appalachian Basin (e.g., Dennison, 1977; Donaldson and Shumaker, 1981; Beardsley and Cable, 1983; Shumaker, 1986b, 1996; Thomas, 1990). In controlling sedimentation, moreover, these structures also affected the later distribution of reservoirs, source beds and sealing units in the basin (e.g., Beardsley and Cable, 1983; Sanford et al., 1985; Harper, 1989; Shumaker, 1996). Other Grenvillian parts of the Laurentian margin were apparently rifted away during the breakup (Figure 3), stranded in oceanic crust, translated elsewhere along strike, and later reaccreted (e.g., Faill, 1997a; Waldron and van Staal, 2001; Bartholomew and Tollo, 2004; Hatcher et al., 2004; van Staal, 2005; Schoonmaker and Kidd, 2006) as part of the supracrustal load that would help generate Appalachian foreland basins. Many of these reaccreted basement fragments and ribbon-like microcontinents have undergone intense deformation during one or more orogenies and are included as internal massifs in the Appalachian internides.

4. Paleogeographic/Paleoclimatic Framework Climate not only influences the tectonic style of adjacent orogens (Hoffman and Grotzinger, 1993), but along with eustasy and tectonics, is a major allocyclic control on foreland-basin sediment supply and sedimentation (Beerbower, 1964; Flemings and Jordan, 1989; Jordan, 1995; Cecil, 2003). Moreover, as regional climate is largely controlled by global, zonal climate belts through which continents migrate and by geographic features like mountain belts and high plateaus that form rain shadows, paleogeography through time may also dictate the sedimentary history of basins (Lees, 1975; Heckel and Witzke, 1979; Ettensohn et al., 2002a; Cecil and DuLong, 2003; Cecil et al., 2004). From the Late Precambrian breakout of Laurentia through Early Cambrian time, the Appalachian area migrated from about lat 601 to 401 S and was largely characterized by siliciclastic sedimentation, and from Middle Cambrian to Late Mississippian time, the Appalachian area passed through various latitudes ranging from 351 to 151 S, where carbonates and evaporites typically predominated (e.g., Scotese, 2003). By latest Mississippian time, the area began moving northward into the southern tropical belt within 51 of the equator, and by the end of

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Figure 4 Paleozoic geologic time scale, showing the occurrence and relative duration of synrift, postrift passive margin, and 13 third-order, tectophase cycles (numbered) in the Appalachian Basin as a relative sea-level curve, compared with generalized sea-level curve (modi¢ed from Ross and Ross, 1988; Read, 1989; and Dennison, 1989), trends in long-term paleoclimate and Southern-Hemisphere ice volume (glaciation) in four respective vertical sections. Note correlation between major glaciation and periods of humidity. Relative sea-level curve in the ¢rst section (except for Sauk Sequence; after Read, 1989) is based on facies and facies distribution in the Appalachian Basin, and the second-order and many third-order cycles compare well with generalized cycles from the second column based on other criteria. The third-order, Middle Ordovician--Pennsylvanian, unconformitybound, transgressive--regressive cycles in the ¢rst section largely re£ect relative Appalachian sea-level responses to tectonism (tectophases), whereas relative sea-level responses in Cambro-Ordovician Sauk Sequence (¢ner stipple) probably largely re£ect thermal and eustatic mechanisms. Columns immediately to the right of the sea-level curve in the ¢rst section represent respective Appalachian orogenies, global (?) orogenic cycles, and Sloss (1963) sequences. The unusual length of Mississippian--Early Pennsylvanian, cycle 11 may re£ect circumstances in the Late Acadian (Neo-Acadian) accretion of Carolina and/or some coincidence with Ouachita or Alleghanian orogenies. Sloss sequences were de¢ned by major regional, and perhaps global, unconformities that are also present in much of the Appalachian Basin; most also coincide with the inception of major Appalachian orogenies. Using the names of Dennison (1989), these unconformities are labeled on the sea-level curve: L, Lipalian; O, Owl Creek; C, Cherokee; W,Wallbridge; and M, Monday Creek. The Siluro-Devonian, Helderberg interval (H) in the ¢rst section probably represents a transitional, orogen-collapse phase separating the larger Caledonian and Variscan--Hercynian orogenic cycles, and two minor third-order cycles within it may re£ect convergence of the Carolina terrane, or could be associated with either of the orogenies that bound it. The superimposed dotted curve re£ects Cambrian--Early Ordovician thermal subsidence and the two second-order orogenetic cycles. Adapted from Frakes et al. (1992), Cecil et al. (2004), and Ettensohn (2005); ages on left (Ma) from Gradstein et al. (2004a, 2004b).

Permian time, was within the northern tropical belt between the paleoequator and lat 51 N, where siliciclastic sedimentation again dominated. This generalized clastic-dominated (Neoproterozoic–Early Cambrian) to carbonate-dominated (Early Cambrian–Middle Mississippian), and return to clastic-dominated (Late Mississippian–Permian) pattern of sedimentation during the history of the Appalachian Basin was apparently controlled in part by zonal climate changes during northward movement of Laurentia/Laurussia through the paleolatitudes (Scotese, 2003; Cecil et al., 2004). During Late Precambrian to Early Cambrian time, the basin was situated in the humid to subhumid, low-pressure belt of westerlies, where clastic sedimentation typically predominates because of increased moisture. From Middle Cambrian to latest Mississippian time, in contrast, the basin area moved into the high-pressure, arid, subtropical belt of easterly trade winds, where aridity and decreased

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moisture favor carbonate and evaporite deposition. By latest Mississippian through Permian time, the basin moved into the low-pressure, humid to perhumid, tropical belt of the equatorial region, where increased moisture facilitates weathering as well as the transport and deposition of clastic sediments. Although zonal atmospheric circulation belts no doubt contributed to large-scale patterns of sedimentation in the basin (e.g., Lees, 1975; Heckel and Witzke, 1979; Cecil and DuLong, 2003; Cecil et al., 2004), some of the more locally occurring lithologies for which the basin is famous, like pure orthoquartzitic sandstones, iron-rich sediments, evaporites, cherts and black shales, may owe their origins to more complex interactions between climate, eustasy and tectonics (e.g., Ettensohn, 1985a, 1985b, 1994; Cecil et al., 2004). Other climatic factors represented in the basin include oscillations related to glacial–interglacial cycles in icehouse-to-greenhouse transitions (Fischer, 1984; Cecil et al., 2004). Although continental glaciation probably existed for at least 90 percent of Phanerozoic time somewhere on Earth (Matthews, 1987), three major periods of glaciation during Late Ordovician–Silurian, Devonian–Mississippian and Pennsylvanian–Permian times (Figure 4) appear to have been especially significant in controlling sediment distribution, unconformities and cyclicity in the Appalachian Basin (Ettensohn, 1994; Cecil, 1990; Cecil et al., 2004). Possible Late Devonian tillites may even be present in east-central parts of the Appalachian Basin (Cecil et al., 2004).

5. Eustatic Framework Several workers have championed the predominance of eustasy in explaining important aspects of Appalachian Basin stratigraphy (e.g., Dennison and Head, 1975; Johnson et al., 1985; Boswell and Donaldson, 1988; Filer, 1994; Brett et al., 1998). However, if the Appalachian Basin is in large part a composite foreland basin generated by deformational loading during four adjacent and nearly continuous, craton-margin orogenies, tectonism in the form of flexural subsidence and uplift and in the reactivation of various basement structures must have been an influence so overwhelming at times as to muffle the effects of coeval eustasy. Nonetheless, if continental glaciation was present to some extent during most of Phanerozoic time (Matthews, 1987) and persistent plate movement sustained tectonoeustasy (Hays and Pittman, 1973), some form of eustatic influence also must have been present. Perhaps the best way to estimate the relative influence of various sea-level controls is to approximate recurrence intervals from the stratigraphic cyclicity attributable to them (Dickinson et al., 1994). Larger scale cycles, for example, in the range of 106–108 years (2nd–4th order), most likely reflect tectonic influence or the interplay of flexural and eustatic processes of similar rate, whereas smaller scale cycles on the order of 104–105 years (5th–6th order) exceed flexural rates and probably record high-frequency eustatic or climatic oscillations (Dickinson et al., 1994; Goodman and Brett, 1994). Movements on individual structures may induce comparable, small-scale changes at similar rates, but may be distinguished because of their very local influence. Hence, by approximating recurrence intervals and referring to relative sea-level curves developed through analysis of sections both within and beyond the Appalachian Basin (e.g., Dennison and Head, 1975; Vail et al., 1977; Johnson et al., 1985; Ross and Ross, 1987, 1988, 1996; Dennison, 1989; Johnson, 1996), especially during the Early Paleozoic drift phase and during periods of major glaciation, eustatic signals that were important to Appalachian sedimentation should be conspicuous. Figure 4 shows second (107–108 years)- and third-order (106–107) sea-level cycles interpreted from the Appalachian Basin, which are largely inferred to reflect nearby tectonic events, compared with a relative, Paleozoic, sea-level curve derived from the work of Ross and Ross (1988, 1996) and Dennison (1989).

6. Flexural Modeling of Foreland-Basin Sedimentation The Appalachian Basin is an extensive area, containing a very diverse series of rocks in many different named units, ranging in age from Neoproterozoic to Early Permian (?). After the Early–Middle Ordovician transition, many basin rocks occur in repetitive, third-order, unconformity-bound sequences (Figure 4) that commonly contain a similar succession of lithologies. The Early–Middle Ordovician transition is commonly represented by a major unconformity, which has been interpreted to represent the initiation of convergence on the new Laurentian margin and ongoing closure of the Iapetus Ocean. During the rest of Paleozoic time, the Laurentian margin was involved in nearly continuous convergence through four orogenies until final Pennsylvanian–Permian collision formed Pangea. Each of these orogenies produced orogenic highlands that acted as sediment sources and also generated accommodation space in the form of a foreland basin due to the isostatic effects of flexural, deformational loading (e.g., Walcott, 1970; Price, 1973; Beaumont, 1981). This loading generated or relaxed stresses that caused adjacent parts of the craton to rise or subside in long-wavelength

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deformations, which influenced relative sea-level fluctuations and cratonic sedimentary sequences across broad areas. In fact, the broad Cincinnati Arch, which defines the western margin of the Appalachian Basin (Figure 1), may have originated through uplift and inversion of an underlying Precambrian rift basin due to nearly 200 Ma of successive loading in the Appalachian orogen (Shrake et al., 1991; Rast and Goodman, 1994). Nonetheless, Ettensohn (1985a, 1991, 1994, 2005) has shown that each of the unconformity-bound sedimentary sequences mentioned above corresponds in time with tectophases in the four orogenies, and that the repetitive sedimentary content of these sequences can be related to distinct stages of loading and relaxation that normally comprise each tectophase, based on the flexural models of Quinlan and Beaumont (1984), Beaumont et al. (1987, 1988) and Jamieson and Beaumont (1988). Because most Appalachian Basin sequences are lithologically similar and correspond in space and time to specific tectonic phases or events (tectophases) at the adjacent continental margin, the flexural models that explain them become very useful in understanding the sedimentary infill of the basin. A summary of the models, extracted from Ettensohn (1991, 1994, 2004, 2005), follows.

6.1. Cycle origins The typical foreland-basin sedimentary cycle consists of seven parts produced as surface (nappes, thrusts and folds) and subsurface (buried, obducted blocks and flakes) deformational loads accumulate on the cratonic margin, severely loading the adjacent lithosphere. To isostatically compensate for the load, the adjacent lithosphere deforms into a downwarped flexural or retroarc foreland basin just cratonward of the deforming orogen and an uplifted peripheral bulge on the cratonward margin of the basin (Figure 5A). The stacking of thrusts and folds supplies most of the load, but a subordinate component is also produced by sediment loading (Beaumont, 1981; Tankard, 1986). Catuneanu et al. (1997) have also indicated that viscous corner flow in the upper mantle associated with a nearby subducting slab creates a sublithospheric dynamic load that can generate longwavelength subsidence which is effective at distances in the foreland far greater than those typically associated with normal surface and subsurface deformational loads. This loading component may generate additional subsidence across the foreland and mask the presence of the forebulge, but it is more prominent during rapid subduction with shallowly dipping slabs (see Catuneanu et al., 1997). As orogeny proceeds and thrust loads shift cratonward, the foreland basin and peripheral bulge also shift cratonward away from the advancing load. Most of the loading and accompanying basin-and-bulge migration will

Figure 5 Sequential, schematic diagrams showing relationships between foreland-basin generation, sediment in¢ll, and deformational loading. (A) Basin-bulge formation and migration with subsurface loading and little clastic in£ux during the ¢rst three parts of a typical cycle. (B) Part four of a typical cycle showing ‘‘loading-type’’ relaxation resulting from a static, sur¢cial load with a developed drainage net that supplies coarser clastic sediments to the subsiding basin (see Figure 6). The pycnocline is a zone of thermohaline density gradation in deeper basins with decreasing O2 content. Dark stipple, dark, organic-rich muds; large stipple, coarse clastic sediments; wavy lines, unconformities. Adapted from Ettensohn (1997).

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Figure 6 Generalized tectophase cycle, showing sequence of £exural events, typical sequence of lithologies, and transgressive--regressive curve for early subduction-type orogenies. Shallow-water deposits above the unconformity are usually very thin or condensed; deeper water, dark muds, however, are typically thicker and are deposited during the most active phases of convergence that generate rapid subsidence in the foreland basin. Sequences are unconformity-bound, but may be incomplete or eroded from the top (adapted from Ettensohn, 1994).

advance cratonward in a direction nearly perpendicular to the strike of the orogen, but if orogeny is diachronous along its length, basin-and-bulge migration will also shift parallel to the strike of the orogen in time (Ettensohn, 1985a, 1987, 1998). Once active loading ceases, the lithosphere responds by relaxing stress through a series of stages that ultimately give rise to ‘‘post-orogenic’’ clastic wedges (Ettensohn, 2004). Inasmuch as each orogeny typically progresses in a series of pulses or tectophases (Jamieson and Beaumont, 1988), the complete loading– relaxation cycle of each tectophase generates an idealized cycle of lithofacies (Figure 6) over the course of several millions to a few tens of millions (106–107) of years. These durations are in line with recent dating and isotope work that suggests the likelihood of short-lived orogenic cycles (o13 Ma), including even shorter metamorphic events, during subduction tectonics (Camacho et al., 2005). The major parts of each cycle and their likely origins (Figure 6) are briefly discussed below. 6.1.1. Basal unconformity The initial result of loading is bulge moveout and uplift, which results in diachronous uplift of the foreland (Quinlan and Beaumont, 1984) (Figure 5A). The resulting bulge will generally be one to two times the width of the adjacent foreland basin, and unconformity formation proceeds from the foreland basin cratonward due to subaerial exposure or elevation into wave base on the migrating bulge (Figure 5A). Because unconformity formation migrates cratonward toward areas of decreasing subsidence and deposition on or near intracratonic highs, these unconformities typically ‘‘open up’’ cratonward, such that the space–time value of the unconformity lacuna increases in that direction. The distribution of the resulting unconformity is generally localized to parts of

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the foreland basin and adjacent craton near the locus of tectonism and is approximately parallel to the strike of the associated orogen (Ettensohn, 1993, 1994). In the larger Appalachian orogen, where much of the tectonism was localized at continental promontories that were subject to greater shortening and resulting deformation (Dewey and Burke, 1974; Dewey and Kidd, 1974; Ettensohn, 1994), the distribution of unconformities is generally asymmetric toward the involved promontories, even within the foreland basin. However, in parts of the foreland basin most proximal to those promontories, unconformities may not develop, because deformational loading is so immense and persistent at the promontory that resulting subsidence in the adjacent foreland-basin depocenter readily offsets any effects of bulge uplift. Thus, unconformity formation in time and space can be an important indicator of tectonic influence (Ettensohn, 1991, 1993, 1994; Ettensohn and Pashin, 1997). Even during major tectonism, eustatic components may also be present and are commonly suggested by presence of a bathtub-ringor bull’s-eye-like pattern of conformable strata that interrupts normal unconformity distribution near the former basin center (Ettensohn, 2007). The unconformities that define the initiation of tectophases are typically regional in nature (Ettensohn, 1993, 1994), but some of the same unconformities developed into the much more extensive, continent-wide, or possibly global, surfaces that Sloss (1963) used to define his sequences (Figure 4). Such unconformities generally seem to record the initiation of significant new episodes of continent-scale tectonism that mark brief periods of major plate reorganization. At these times, the response of the continent to new subduction or collision seems to have been one of impedance, resulting in broad, upward lithospheric deflections across large parts of the continent during few-million-year episodes, which Sloss and Speed (1974) called ‘‘emergent.’’ Dickinson (1974, p. 22) called such a response ‘‘braking’’ inasmuch as subduction is initially retarded or ‘‘braked’’ by the immobile continent being subducted. 6.1.2. Shallow-water transgressive deposits After bulge moveout, deformational loading leads to rapid subsidence of the foreland basin, and shallow transgressive seas initially move across the unconformity. Subsidence is typically so rapid that the resulting shallow-water deposits are very thin or condensed (Figure 6). Nonetheless, in subtropical areas with little previous clastic influx, transgressive carbonates typically develop, and in areas that just experienced tectonism, residual clastic debris on the surface, sometimes exposed long enough to become reworked by eolian processes (Grabau, 1932, 1940; Cecil et al., 1991), will generate very mature, thin, shoreface sand deposits in this position. Whether clastic or carbonate, rapid subsidence and consequential deepening ensure that such deposits are short-lived. The situation is different in basins that form adjacent or peripheral to the suture belt and are on the plate margin being consumed. These basins are called peripheral foreland basins by Dickinson (1974), and in the Appalachian setting, commonly formed in the large structural reentrants between promontories. Because such margins are being rapidly drawn toward the subduction zone, the peripheral basin that forms, as marginal crust is faulted and depressed, will initially receive thick, flysch-like, clastic sequences from the adjacent arc and suture zone immediately on top of the unconformity (see Lavoie, Chapter 3, for examples from the Quebec reentrant). Later, if subduction vergence changes, these basins may be destroyed or tectonically collapsed beyond recognition as parts of the hinterland internides. 6.1.3. Dark-mud sedimentation Subsidence in the foreland basin is largely an isostatic response to deformational loading caused by crustal shortening and load transfer along a steep basement ramp. Early in the development of an orogen, it is likely that such ramps are only able to accommodate up to 20 km of vertical deformation without creating major subaerial topography or source areas (Jamieson and Beaumont, 1988), so that initially much of the deformation must occur in the subsurface or in subaqueous environments that generate little or no subaerial relief (Karner and Watts, 1983) (Figure 5A). As a result, no major source of externally derived sediment is available to begin basin infilling. Hence, after shallow-water deposits are drowned, the foreland basin experiences sediment starvation, and in the absence of major clastic influx, organic mater from the water column and suspended silts comprise the predominant sediment input. At the same time, the basin continues to undergo rapid subsidence with which the scant sedimentation cannot keep pace. Consequently, the deepening water column becomes stratified, and the organic matter is buried and preserved as dark or black muds in resulting oxygen-deprived (dysoxic or anoxic) environments. Based on modern analogs (Rhoads and Morse, 1971; Byers, 1977; Ettensohn and Barron, 1981), stratification may ensue once water in the basin attains depths greater than 150 m. Although dark muds typically characterize rapidly subsiding proximal and central parts of a foreland basin, at the same time in more distal parts of the basin and on adjacent parts of the craton, reduced subsidence typically produces coeval, transgressive deposits of carbonates or light-colored shale overlying the basal unconformity. Nonetheless, the black or dark

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shales are the most distinctive part of each sedimentary cycle (Figure 6), and tracking their successive distributions in time and space is probably the best way to relate the cyclic sedimentary responses of foreland basins (Figure 4) to tectonic cues (Figures 7–10) (Ettensohn, 1994, 1998, 2005). The intense deformational loading and concomitant foreland-basin subsidence at this time clearly reflect ongoing plate convergence and subduction. However, yet another manifestation of subduction at this time should be arc volcanism, and although volcanics are very rare in foreland basins, the presence of a volcanic arc may be reflected in foreland-basin sedimentation through beds of wind-blown ash, preserved as bentonites. Although bentonites may occur in any part of a foreland-basin sequence, they are typically more common in the shallowwater, transgressive and early dark-mud, depositional phases of a foreland-basin cycle, because during these phases of deposition, subduction is most active.

Figure 7 Schematic map of the Appalachian Basin area showing the relative positions of Taconian tectonic highlands, successive, dark-shale, foreland basins, continental promontories, and the relative positions peri-Gondwanan microcontinents during Late Ordovician--Early Silurian time. Extent of successive Taconian foreland basins (Sevier, Martinsburg, and Power Glen-Cabot Head) was determined by mapping the distribution of respective lower black-shale units in each cycle, and the northwestward migration of these foreland basins in time and the their asymmetry toward di¡erent promontories shows migration and focus of the successive tectophases that generated them. Each basin approximates the maximum aerial distribution of the respective tectophase sequence. The Sevier basin was the product of the Blountian tectophase; the Martinsburg basin was the product of the Taconic tectophase; and the Power Glen-Cabot Head was the product of Early Silurian tectophase 3. Figure 4 shows the temporal, tectonic, eustatic, and climatic particulars of the related,Taconian, tectophase cycles, whereas lettered positions indicate locations of the sections in Figures 8--10, which show the temporal and lateral extent of black-shale units and parts of the Taconian tectophase cycles. CT ¼ Carolina terrane; GA ¼ Gander terrane; AV ¼Avalon terrane. Adapted from Ettensohn (1991, 1994).

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Figure 8 Schematic, southwest-northeast, Middle--Upper Ordovician section (section ABFC in Figure 7), paralleling the strike of the Appalachian Basin and showing repetition of foreland-basin tectophase cycles for the Blountian and Taconic tectophases. Each cycle begins with an unconformity, a thin transgressive carbonate and a major black-shale unit. Note the northeastward migration of tectophase sequences in time and of the Martinsburg-Utica dark shales, which re£ect the diachronous northeastward migration of Taconian tectonism on the southeastern margin of Laurentia. The post-Sauk unconformity is the same as the Owl Creek unconformity in Figure 4. No vertical scale intended. Adapted from Ettensohn (1991, 1994).

6.1.4. Flysch-like sedimentation Dark-mud deposition occurs throughout central parts of the foreland basin as long as active orogeny and deformational loading continue, but once active thrust movement declines and tectonic quiescence ensues, the deformational load becomes static. The lithosphere responds to the now static load by relaxing stress, as a result of which the proximal foreland basin subsides while the peripheral bulge is uplifted and shifts toward the subsiding load (Figure 11A). By this time, substantial subaerial relief has been generated by emplacement of a surficial load (fold-thrust belt), and surface drainage nets have had time to develop (Figure 5B). As a result, coarser grained clastic debris is eroded and transported into the foreland basin in the form of deeper water deltaic deposits, turbidites, contourites and debris flows (Figures 5B and 6), while some nearshore clastic sediments may be redistributed as tempestites by storms. Short-system sedimentation (Fichter and Poche´, 1993) like this that involves relatively rapid transportation to the basin and largely contains darker, deeper water, commonly immature, siliciclastic components, is herein referred to as flysch-like. While the subsiding foreland basin is being filled with deeper water, flysch-like sediments, the adjacent bulge is uplifted and migrates basinward (Figure 11A), generating a regional unconformity that may truncate previously deposited flysch-like sediments and/or a regressive carbonate sequence atop the bulge, depending upon the relative disposition of sea level. Ettensohn (1994) has called the flexural process at this stage ‘‘loading-type relaxation,’’ and it results in a regressive, forelandbasin infill (Figures 5B and 6). In subduction-type orogenies, the foreland or retroarc basin on stationary lithosphere is commonly ‘‘balanced’’ as a sediment sink by a peripheral, forearc or pro-foreland basin on converging lithosphere (Dickinson, 1974; Willett et al., 1993; Johnson and Beaumont, 1995). Such ‘‘double-sided’’ or ‘‘symmetrical’’ orogens theoretically assure the presence of two ‘‘sinks’’ for the dispersal of debris eroded from the subaerial load. Where one of these sinks is absent, as for example in the loss of basins on converging crust in some collisional orogens, or where climatic factors, which control weathering and erosion, are focused more on one side of an orogen than on the other, the other sediment sink or basin may fill more rapidly and aggrade upward into marginal-marine or terrestrial settings that would not normally be expected at this stage. 6.1.5. Thin, regressive, shallow-water, carbonate or shale blanket As the adjacent, static, surface load is eroded, the rate of siliciclastic influx eventually exceeds basin subsidence rates, and the foreland basin fills or overflows with siliciclastic sediment. This infilling of the foreland basin,

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Figure 9 Schematic northwest-southeast section (section EH on Figure 7), nearly perpendicular to the strike of the central Appalachian foreland basin, showing the cratonward migration of two Taconian foreland-basin tectophase cycles in time. Note occurrence of dark shales and other parts of tectophase cycles. The breaks at the bottom and top of the section re£ect the Owl Creek and Cherokee unconformities, respectively (see Figure 4). No vertical scale intended. Adapted from Ettensohn (1991).

combined with greatly lowered source areas and a waning supply of clastic sediment, may set the stage for deposition of an extensive, but thin, blanket of shallow-water carbonates or mixed carbonates and shales (Figure 6). If the basin occurs at arid, subtropical latitudes, carbonates typically predominate; however, if the basin is present in the rainy, equatorial or temperate latitudes, shales or mixed carbonates and shales are more likely. These blanket carbonates or shales mark the culmination of shallowing and regression that began with the cessation of loading and the beginning of loading-type relaxation (Figures 6 and 11B). The carbonate or shale blanket may become very widespread at this time because the filled basin and lowered source areas briefly approach the same elevation, enabling shallow seas to spread widely. 6.1.6. Thin, transgressive, marine, precursor sequence The above phase of ‘‘elevational equilibrium’’ is short-lived, because the area of the former orogen and foreland basin begin to rebound upward in isostatic response to the lost load (Figure 11B). During this ‘‘unloading-type relaxation,’’ previously eroded upland areas and adjacent parts of the foreland basin rebound and a compensating ‘‘anti-peripheral bulge,’’ called a peripheral sag by Ettensohn (1994), forms and moves toward the rebounding area (Beaumont et al., 1988) (Figure 11B). A short-lived, transgressive sequence of open-marine shales, or shales and carbonates, is deposited in the peripheral sag (Figure 6), but the sequence is really only a thin, marine precursor to the thick wedge of prograding siliciclastic sediments that will follow. 6.1.7. Marginal-marine and terrestrial clastic wedge The final part of the cycle is characterized by a cratonward-prograding wedge of predominantly marginalmarine-to-terrestrial, siliciclastic sediments, which have been described as ‘‘post-orogenic,’’ ‘‘molasse-like’’ or ‘‘deltaic’’ (Figure 5). Because the rebounding load consists of formerly beveled highlands and previously deposited

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Figure 10 Schematic west-east section, nearly perpendicular to the strike of the northern Appalachian foreland basin (section GF on Figure 7), showing the nature and disposition of the three Taconian tectophases. The Black River carbonates, though unconformity-bound, re£ect an atypical and very poorly developed tectophase sequence, probably re£ecting the great distance from the Virginia locus of the Blountian tectophase. Note the cratonward migration of the ¢nal two tectophases, as well as the cratonward migration and ‘‘younging’’of the Utica Shale. Unconformities below the Utica apparently represent reactivation of local structures; the post-Sauk unconformity is the same as the Owl Creek unconformity in Figure 4. Symbols as in Figure 9; no vertical scale intended (from Ettensohn and Brett, 2002).

Figure 11 Schematic diagram showing two types of £exural response to lithospheric stress relaxation in the relaxational part of a tectophase cycle (redrawn from Beaumont et al., 1988). (A) ‘‘Loading-type’’ relaxation --- thrust migration ceases and the resulting static load causes lithospheric relaxation with major basin subsidence in part 4 of a typical cycle. (B) ‘‘Unloading-type’’ relaxation --- erosional unloading results in rebound and erosion near unloading area in parts 5--7 of a typical cycle; see Figure 6 for the complete cycle.

foreland-basin sediments (Figure 11B), sediments eroded from the rebounded area will be primarily fine-grained, composed mainly of siltstone, silty shale, shale, mudstone or shaly carbonate (Ettensohn, 1994, 2004, 2005). Because the flexural process begins from a state of approximate elevational equilibrium at or near sea level, a single cratonward-dipping paleoslope becomes established (Figure 11B). Because of low gradients, proximity to the sea, and poorly differentiated energy gradients, local features and autogenic sedimentary variations may generate many vertically and laterally changing depositional environments in this setting; hence, a facies mosaic including redbeds, paleosols, coals and even thin carbonates, in close association with siliciclastics, may be characteristic (Figure 6). Redbeds are especially common, and because of the common association of redbeds

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Figure 12 Highly schematic, composite, west-east section from north-central Ohio to east-central NewYork in the northern Appalachian Basin. Parts of four Acadian tectophase sequences are shown, but only the ¢nal Mississippian sequence went to completion, and it is not completely shown here. Each of the early three tectophases apparently did not go to completion before the next one started, and erosion from the succeeding tectophase partially destroyed the sedimentary record of the previous one. Note the cratonward migration of succeeding tectophase sequences and the fact that the third tectophase is composed of subcycles that may re£ect the movement of individual thrust complexes. No scales intended (from Ettensohn, 1987, 1994, 2005).

with the terrestrial parts of deltas, many Appalachian Basin redbed units in this part of the cycle, like the Queenston (e.g., Figures 8–10), Bloomsburg, Catskill, Bedford-Berea and Pennington-Mauch Chunk, have been called ‘‘deltas.’’ In reality, they are delta complexes or tectonic delta complexes, which may reflect several different marginal-marine environments (Friedman and Johnson, 1966; Ettensohn, 2004) that prograded far beyond the foreland basin, giving the appearance that the foreland-basin has ‘‘overflowed’’ onto the craton. This situation may be similar to what Jordan (1995) has called an ‘‘overfilled’’ foreland basin, except that in the models used herein, overflow or overfilling are commonly typical only in the later stages of relaxation and are controlled more by flexural processes than by surficial or climatic processes. During this relaxation process, proximal parts of the foreland basin may be cannibalized and generate an unconformity that ‘‘opens up’’ toward the tectonic highlands (Goodman and Brett, 1994). Moreover, any erosion that accompanies the bulge-moveout phase of the next tectophase or orogeny will readily subsume such an unconformity and parts of the underlying succession. Hence, an incomplete cycle of lithologies (e.g., Figures 10 and 12) may reflect erosion during the succeeding tectophase or the advent of a new tectophase before the sedimentary expression of the previous one was complete (Ettensohn, 1994, 2004). 6.1.8. Model summary and implications As Figure 6 shows, the typical, tectonic, foreland-basin cycle consists of seven parts in ascending order: (1) basal unconformity; (2) shallow transgressive carbonates or shoreface sands; (3) dark shales; (4) a flysch-like clastic sequence; (5) a thin blanket of shallow-water carbonate or carbonate and shale; (6) a thin, open-marine, shale sequence and (7) a thick, marginal-marine, clastic sequence with redbeds. During deposition of parts (2) and (3) of the cycle, it is more likely that the foreland basin will exhibit two paleoslopes and be underfilled, whereas

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during deposition of part (7), only one paleoslope is likely and the basin is more likely to experience overfilling (Figure 11). Because orogenies occur in pulses or tectophases (Johnson, 1971; Jamieson and Beaumont, 1988), one or more of these cycles will constitute the sedimentary record of an orogeny in its foreland basin (Figure 4). As an example, the cyclic sedimentary record of two or three tectophases from the Ordovician–Silurian Taconian orogeny in the Appalachian Basin (Figures 4 and 7) is shown in a section parallel to basin strike (Figure 8) and in two sections perpendicular to strike (Figures 9 and 10). Individual cycles, however, may not exhibit a complete sequence of lithologies because of erosion at the base of the succeeding cycle (Figures 10 and 12), because the succeeding tectophase began before the sedimentary expression of the previous one was complete (Figure 10, tectophase 3), or because the area of the section during deposition was too distant from the locus of major tectonism and resulting subsidence (Figure 10, Blountian tectophase). It is also possible for subcycles of blackshale and flysch-like clastics to repeat several times in a single tectophase (Figure 12), each sub-cycle perhaps reflecting movement of major thrust systems during the tectophase. Based on current observations, the simplest noted cycles consist of an unconformity, dark shales and a flysch-like sequence (Figure 10, tectophase 3). However, one consistent feature of any foreland-basin cycle or sub-cycle is that the sedimentary record of each successive tectophase migrates farther cratonward (Figures 7–10, 12), or farther along strike during oblique convergence or in transpressive orogenies (Figures 7 and 8) than did the previous one (Ettensohn, 1985a, 1991, 1994, 1998). The best way to observe this progressive, perpendicular-to-strike or parallel-to-strike movement of the cycles is to map the distribution of the dark-shale part of each cycle (Ettensohn, 1985a, 1994, 1998) (Figure 7); not only do the dark shales represent the time of maximum subsidence and transgression in each cycle, and hence its greatest extent, but they are also probably the most distinctive and easily recognized of cycle lithologies, even in cores and geophysical logs. The development of such cyclic sequences in foreland basins (Figure 4) seems to be most characteristic of orogenies with some subduction component during early parts of the convergence history of the adjacent continental margin. During early parts of convergence, mass transfer from one plate to another and intra-plate shortening mainly occur along a steep basement ramp that defines the rifted continent margin. Much of the transferred load, moreover, would have been concentrated narrowly on or near the basement ramp, and the isostatically compensating foreland basin would have been correspondingly narrow and deep. However, in successive tectophases, it is likely that the deformational front would have surmounted the continental margin and developed sufficient topography to advance substantially onto the craton as a surficial load (Jamieson and Beaumont, 1988). Because the deformational load is now more expansive and spreads itself across a greater area, the resulting lithospheric flexure generates a broader, shallower foreland basin (Karner and Watts, 1983), and the successive broadening of foreland basins during any one orogeny is a common pattern. During the Taconian orogeny, for example, not only did successive foreland basins move cratonward, reflecting the advancing load, but they became greater in area (Figure 7), and the contained cyclic sequences, especially dark-shale and flysch-like parts of the cycles, overall became shallower with time based on various sedimentary criteria. In fact, except for shale color, in the later-formed foreland basins of an orogeny, foreland-basin cycles may become diffuse and difficult to distinguish from normal cratonic sedimentation, pointing out the fact that the same flexural stresses controlling foreland-basin sedimentation may also influence patterns of cratonic sedimentation some distance from the foreland basin. In addition, if a continental margin like the Appalachian margin experienced multiple orogenies, patterns of foreland-basin shallowing and broadening not only continue, but may be amplified by the increasing flexural rigidity of the lithosphere with time (Watts et al., 1982). By the latest of such orogenies when the deformational front has advanced far onto the craton, the resulting foreland basin may be so shallow and distant from the sea that typical cycles do not develop, marginal-marine and terrestrial clastic sediments predominate, and these sediments overflow from the foreland basin to form a siliciclastic blanket that can spread hundreds of kilometers across large parts of the adjacent craton (Ettensohn, 1994, 2004, 2005); however, Flemings and Jordan (1989) have indicated that climatic variations and low thrust velocities may create conditions in the foreland basin that can also generate overflow. The above effects are also enhanced by the fact that lithospheric rigidity gradually increases in time and with each successive orogenic event so that the Appalachian Basin should have been characterized by decreasing rates of subsidence during Paleozoic time, and the general predominance of major foreland-basin facies through time support this model (Tankard, 1986; Willard and Klein, 1990). In particular, during early basin history in Ordovician time, black shales and turbidites predominate; during Devonian–Mississippian time, turbidites and alluvial-deltaic sediments are most abundant; whereas during late basin history in Pennsylvanian–Permian time, marginal-marine and fluvio-deltaic sedimentation predominated. Nonetheless, even during the later parts of basin history, tectophase cycles still appear to be present, but their manifestation is very different than that shown in Figure 6 (Ettensohn, 2005). Finally, there seems to be no evidence from the Appalachian Basin supporting the idea of reciprocal sedimentation (Catuneanu et al., 1997), the idea that deepening in the foredeep was contemporaneous with regression on the forebulge, or vice versa. Although facies certainly differ among distal, forebulge and proximal,

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foredeep areas in all the tectophase cycles, base-level changes seem to have been uniform across both proximal and distal settings. This uniformity may reflect the ineffectiveness of dynamic loading in the Appalachian foreland because of steep subducting slab dips or because the additional, intervening belt of accreting microcontinents that was almost always present in Appalachian orogenies between the subduction zone and the deformed crustal margin meant that the effects of dynamic loading were never transferred as far cratonward as distal parts of the foreland basin.

7. Generating the Appalachian Margin: Late Precambrian–Early Cambrian Rifting and Rift Fill (B765–B535 Ma) The initial Appalachian margin was produced by the Late Neoproterozoic rifting of Rodinia and the breakaway of Laurentia from western Gondwana. A series of intracratonic rift basins developed in the continental interior, while aulacogens and marginal, ensialic, shelf basins developed along a subsiding, zigzagging continental margin (Figure 3) with scarp-bounded, mountainous relief (e.g., Schwab et al., 1988). A relatively detailed account of these events has been preserved in rift-related plutonic, metasedimentary and metavolcanic rocks in the Appalachian area from Vermont and adjacent parts of Canada (see Lavoie, Chapter 3) to North Carolina and Tennessee (Figure 13). These rocks unconformably overlie Grenvillian basement and vary in thickness from a feather edge to more than 16,000 m in eastern Tennessee (King et al., 1958; Patchen et al., 1985a); they generally reflect a three-part process: a phase of incomplete rifting at about 765–570 Ma, full rifting from 570 to 535 Ma, and transition to a passive margin (Tollo et al., 2004b; Badger and Sinha, 2004; Novak and Rankin, 2004), although Rast and Kohles (1986) have suggested an even earlier phase of rifting at least 800 Ma old. The early phase of incomplete rifting is reflected by the Mt. Rodgers formation and related rocks in Virginia and North Carolina, and possibly by parts of the Ocoee Supergroup in Georgia, Tennessee and North Carolina (Figure 13) (Rast and Kohles, 1986), where rift-related rocks unconformably overlie the Blue Ridge basement complex. However, more recent work on the Ocoee is now suggesting that the sequence may be entirely or partially Middle Paleozoic in age and not at all rift-related (Unrug and Unrug, 1990; Unrug et al., 2000), although others suggest that the fossil evidence for these ages is incorrect (e.g., Corrie and Kohn, 2007). Nonetheless, plutons and dikes, as well as volcanic, volcaniclastic and sedimentary rocks, in the Mt. Rodgers Formation do most likely reflect rift-related volcanic centers, and in places these rocks are overlain by tilloid diamictites that indicate alpine glaciation (Rankin, 1993; Rankin et al., 1994; Novak and Rankin, 2004). The high-latitude position of the continent at the time would have certainly been conducive to glaciation. In other areas, especially on the eastern side of the Blue ridge, rift-related rocks show a transition from non-marine to deep-marine sedimentation, suggesting deposition in rift basins on thinned, attenuated crust with access to the sea just prior to actual continent separation (Wehr, 1985; Wehr and Glover, 1985). By about 570 Ma, actual separation of Laurentia and inception of the Iapetus, Theic and Rheic ocean basins had begun as a consequence of Gondwana amalgamation (van Staal et al., 1998), and the separation generated a series of intracontinental rifts subparallel to the new Laurentian margin in which volcanics and minor terrestrial clastics accumulated, as well as continental-margin basins, in which marine clastic sediments accumulated (e.g., Faill, 1997a; Gates and Volkert, 2004). Also generated in the final rifting event were a few slivers or ‘‘ribbons’’ of Grenvillian crust, like the Goochland, Baltimore?, Chain Lakes, and Dashwoods terranes or microcontinents, that were rifted away and stranded in oceanic crust some distance from the main margin (Faill, 1997a; Waldron and van Staal, 2001; Bartholomew and Tollo, 2004; Hatcher et al., 2004; van Staal, 2005; Schoonmaker and Kidd, 2006) (Figure 3). This phase of rifting is probably best known for the Catoctin flood basalts of Pennsylvanian, Maryland and Virginia (Figure 3) and related mafic dikes (e.g., Faill, 1997a; Badger and Sinha, 2004), but approximately equivalent rift magmatism is also reported from eastern Canada (Lavoie et al., 2003; van Staal, 2005; Lavoie, Chapter 3). The slightly older age of rift-related dikes in the north suggests that final rifting was diachronous and proceeded from north to south (Faill, 1997a). The growing knowledge of these widely distributed basins and the possibility of more inboard basins beneath Valley and Ridge cover has suggested to some the former presence of a broad, Basin and Range-like rifted margin rather than the narrow margin previously interpreted (Gates and Volkert, 2004). Although no evidence of rifts or rift-fill sediment is apparently preserved in New England, Schwab et al. (1988) have indicated that offshore, continental-rise, sedimentary prisms of mainly Early Paleozoic age may contain Late Neoproterozoic, non-marine, continental-edge, conglomeratic infill in the Pinnacle, Hoosac, Tyson, Underhill and Cavendish formations of western New England. In the Canadian Humber succession of the Quebec reentrant (Figure 3), an allochthonous continental-margin domain thrust westward during the Taconian orogeny, rift-related lavas and coarse-grained graben, slope and rise sediments of Late Neoproterozoic to Early

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Figure 13 Map of the eastern Appalachian area showing the approximate distribution of Grenvillian, Mesoproterozoic crystalline massifs (black) and possible Iapetan, Neoproterozoic rift-basin deposits (stipple). The western line of massifs are the external massifs that approximate the eastern limit of the preserved Appalachian Basin; they probably represent parts of the former Iapetan, rifted margin that were later thrust westward on megathrusts. The more isolated massifs to the east are internal massifs that represent former continental fragments that were stranded in oceanic crust during rifting and re-accreted in subsequent orogenies. C, Catoctin volcanics; M, Mt. Rodgers Group; O, Ocoee Group; dotted lines, eastern U.S. state boundaries (see Figure 3) (distributions from Rankin et al., 1989; Gates and Volkert, 2004).

Cambrian age are present in the Labrador, Curling, Oak Hill, Caldwell Sillery, Saint-Roche and Shickshock groups (see Lavoie, Chapter 3). Yet more distal, Iapetan, intracontinental, crustal extension is represented by a series of yet younger, more interior, rifts that developed during Early to Middle Cambrian time from Alabama in the south to New York, Ontario and Quebec in the north and as far west as western Tennessee and eastern Arkansas (Fisher, 1977; Sanford, 1993b; Shumaker and Wilson, 1996) (Figure 1). Most are only known from subsurface information, and the best known of these rifts is the Rome trough (Figures 1 and 2). Elsewhere in the Quebec Basin, sandstone, conglomeratic and arkosic redbeds of the Covey Hill Formation were deposited in such a rift-related terrane during Late Neoproterozoic and earliest Cambrian time (Sanford, 1993b). Subsidence in the rifts was most active from Early Cambrian to Early Ordovician time (e.g., Ryder et al., 1992) and reactivation of bounding faults influenced sediment thickness, facies distribution, later mineralization and the location of the Appalachian foreland basin through time (e.g., Heyl, 1972; Shumaker, 1986a, 1986b, 1996; Shumaker and Wilson, 1996; Faill, 1997a; Ettensohn et al., 2002b).

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By latest Neoproterozoic to Early Cambrian time, the final transition to a drift-stage or passive continental margin is reflected by a blanket of terrestrial to marginal-marine, quartzose clastic sediments that conformably to unconformably overlie Grenville basement and rift-related rocks across the Appalachian area (e.g., Dennison, 1989; Simpson and Eriksson, 1989; see Lavoie, Chapter 3).

8. The Appalachian Passive Margin By latest Neoproterozoic time, large parts of the eastern Laurentian margin became a stable, passive margin that faced the opening Iapetus Ocean basin. Combination of thermal subsidence and change from Neoproterozoic icehouse to greenhouse conditions provided concomitant sea-level rise and resulting transgression onto the new margin (Figure 4). Initial sediments deposited during the transgression were latest Neoproterozoic to Early Cambrian quartzose clastics derived from the Laurentian craton to the northwest and deposited as a widespread, coastal-plain blanket along the margin (Figure 14). Because of greater subsidence along the margin, the resulting, wedge-shaped body of clastic sediments is older and thicker in eastern parts of the Appalachian Basin, where the rocks are predominantly Early Cambrian in age, but they become younger and thinner in western and northern parts of the basin, where the rocks are predominantly Late Cambrian in age.

Figure 14 Schematic northwest-southeast section across north-central parts of the Appalachian Basin from northernVirginia to northwestern Ohio, showing overall stratigraphic relationships and major stratigraphic units; datum is the base of the Silurian (Tuscarora Ss. or Brass¢eld Lms.). In the uppermost Precambrian to Middle Ordovician section, units are relatively uniform and widespread, re£ecting gentle thermal subsidence in a post-rift, passive-margin setting; hence, major stratigraphic variations re£ect eustasy or local rift-related subsidence on an extensive carbonate platform (e.g., Rome trough). The Cambro-Ordovician carbonates of the Elbrook-Conasauga, Conococheague, Beekmantown, and Knox formations are parts of what has been called the ‘‘Great American Bank.’’ From Late Ordovician to the end of Paleozoic time, the platform was destroyed by tectonism and tectonic di¡erentiation along basement structures led to prominent stratigraphic di¡erentiation and a post-Middle Ordovician facies mosaic. The Waynesboro-Rome, Martinsburg, Salina evaporites, and Devonian black shales formed important Alleghanian de¤collement zones. Abbreviations: Lo, Loudon Fm.; W,Weverton Ss.; MS, Mt. Simon Ss.; K, Kerbel Fm.; E, Eau Claire Fm.; M, Martinsburg Sh.; Br Brass¢eld Lms.; T,Tuscarora Ss.; L, Lockport Dol.; Bl, Bloomsburg Fm.; S, Salina Gr.; Co, Columbus Lms.; O, Onondaga Lms.; Hun, Huntersville Chert; Or, Oriskany Ss.; N, Needmore Sh.; Ma, Marcellus Sh.; Bu, Burket Sh.; Mi, Middlesex Sh.; R, Rhinestreet Sh.; H, Huron Sh.; C, Cleveland Sh.; Oh, Ohio Sh.; Ol, Olentangy Sh.; BB, Bedford-Berea; Su, Sunbury Sh.; Bor, Borden Fm. Formation thickness and scales are approximate. Adapted from Colton (1970) and Shumaker (1996).

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Locally, these transgressive sediments may be gradational with upper parts of the previous rift sequence (Brezinski, 2004), but throughout most of the basin area, they unconformably overlie rift-related strata or Grenvillian basement. Subsidence was largely related to thermal contraction and occurred at rates ranging from a few centimeters to 10 cm/Ka and was concentrated at three hinge lines parallel to the rifted margin that migrated westward in time (Read, 1989). Moreover, varying subsidence rates in three distinct depocenters in Tennessee, Virginia and Vermont, apparently separated by structures related to promontory boundaries, controlled thickness and facies distribution (Read, 1989).

8.1. Latest Neoproterozoic–Early Ordovician pericratonic sedimentation (upper Sauk sequence, B570–472 Ma) After rifting, a wedge of terrigenous to marginal-marine, clastic sediments was deposited along the length of the Appalachian margin in coastal-plain environments that record stabilization of the continent margin (Walker and Driese, 1991). These basal clastic sediments are included in the Weisner Quartzite of Alabama and Georgia, the Chilhowee Group of Tennessee through Virginia to Maryland, the Unicoi Formation of central-western Virginia, the Hardystone–Chickies–Poughquag–Lowerre quartzites of eastern Pennsylvania, New Jersey and eastern New York, the Potsdam Formation of western New York, the Potsdam Group of eastern Canada, including the Nepean Formation in the Ottawa Embayment and the Chaˆteaugnay Formation in the Quebec Basin, the Cheshire and Dalton formations of Vermont and western Massachusetts, and the Mt. Simon Sandstone in the subsurface of the western Appalachian Basin (Poole et al., 1968; Dennison, 1984; Schwab et al., 1988; Sanford, 1993b) (Figure 14). In units like the Chilhowee and Unicoi, the exact location of the Neoproterozoic– Cambrian boundary in the basal part of the section is commonly uncertain (Walker and Driese, 1991), but these basal sandy units always become younger toward the west and northwest at the preserved edges of their distribution. For example, in western New York the Potsdam Sandstone is Late Cambrian in age, but in eastern Ontario and southern Quebec, where it is more than 600 m thick, the Potsdam Sandstone probably spans the Cambrian–Ordovician boundary (Poole et al., 1968; Sanford, 1993b). This basal clastic sequence typically consists of coarse arkose, conglomerate, greywacke, argillite, minor volcanics and orthoquartzitic sandstones and attains maximum thicknesses of about 2,300 m in southwestern Virginia and northeastern Tennessee, but elsewhere more typical maximum thicknesses are in the 300–500 m range, pinching out to a feather edge in western parts of the basin (Colton, 1970; Patchen et al., 1985a, 1985b; Skehan, 1985; Schwab et al., 1988; Read, 1989). In the thickest parts of the Chilhowee clastic sequence, alluvial, meandering-to-braided fluvial, shoreface-to-tidal and shallow-marine environments are represented in succession (Schwab, 1971, 1972; Simpson and Eriksson, 1989), although in the younger, thinner and more cratonward equivalents, shoreface sandstones seem to predominate. Early phases of clastic deposition were apparently controlled by relief and tectonic setting (Walker et al., 1994), whereas later variations in succession appear to largely reflect sea-level changes (Figure 4). In fact, five to six shale-carbonate cycles in Middle and Upper parts of the Cambrian section have been interpreted to reflect shortterm (2–10 Ma), third-order eustatic events superimposed on a longer term eustatic trend (Bond et al., 1988). With continued deepening and shelf stabilization, Lower Cambrian siliciclastic sediments were displaced by a shelf-edge, carbonate bank from Georgia to Vermont represented by the Shady, Tomstown (Figure 14) and Dunham dolostones. These units, ranging from 300 to 1,300 m thick, are the oldest Paleozoic carbonates in the Appalachian Basin area and represent shelf-edge algal reefs and carbonate shoals with cratonward lagoonal and tidal-flat environments (Reinhardt and Wall, 1975; Pfeil and Read, 1980). These carbonates also reflect the initial establishment of a carbonate bank on the trailing edge of a continent, a type of large-scale, tectonic depositional framework first noted by Rodgers (1968). The bank, however, was soon drowned across nearly the entire eastern platform by a late Early Cambrian influx of terrigenous sand and mud (Figure 14, Rome and Waynesboro formations) accompanying regional uplift in the craton interior and further extension across the Rome Trough (Milici and de Witt, 1988; Read, 1989). During the accompanying regression, 150–600 m of shale, sandstone, redbeds and minor dolostone were deposited in a mosaic of shoreface, tidal and very shallow open-marine environments reflected in the Rome, Waynesboro and Monkton formations, and across much of the eastern platform and platform margin, an unconformity developed at the Lower–Middle Cambrian transition (Read, 1989; Faill, 1997a). In the Rome Trough of eastern Kentucky, the subsurface Rome Formation includes alluvialfan and fan-delta deposits related to ongoing extension (Figure 14). Following regression and unconformity development, renewed submergence during Middle and Late Cambrian time resulted in the resumption of carbonate-bank deposition along the shelf edge, reflected in 100–900 m of very shallow subtidal to supratidal dolostones in the Elbrook, Honaker, Stissing, Winooski, and equivalent formations (Figure 14). In western and southern parts of the Appalachian Basin, dolostones

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intertongue with shallow open-marine, lagoonal, or intrashelf-basinal shales from the Conasauga Formation, which apparently reflect prodelta muds shed from the southward-prograding, Kerbel shoal-water delta in central Ohio (Janssens, 1973) (Figure 14). Despite three minor Middle to Late Cambrian regressive pulses of lagoonal or intrashelf basin muds from the delta during deposition of the Conasauga and equivalent dolostones to the east, submergence of the platform continued into Late Cambrian time, and the carbonate bank was re-established across the entire Appalachian region, spreading both westward across the continental interior and eastward as its seaward margin was expanded (Palmer, 1971; Read, 1989). Although renewed deepening was important in re-establishing carbonate sedimentation, as it restricted most clastic influx to shoreline areas, equally important was the continued northward shift of Laurentia relative to the equator. As a result, by Late Cambrian time, the Appalachian area had moved into the arid, evaporative, trade-wind belt (Scotese, 2003) (Figure 4), where carbonate deposition is especially favored. About 50–950 m of dolostone, limestone, and minor siliciclastic sediment were deposited in bank units, which include the Copper Ridge, Conococheague, lower Knox, Gatesburg, Allentown, Whitehall, and Clarendon Springs formations (Figure 14). Many of these units are composed of meter-scale, shoaling-upward cycles that include ribbon carbonates, thrombolites, grainstones, tidal laminites, cryptalgalaminites, and locally, replaced evaporites (Demicco, 1982; Read, 1983). Throughout most of the Appalachian area, this type of sedimentation continued uninterrupted into Early Ordovician time and is represented by parts of the Knox and Beekmantown groups (Figure 14). These dolomitic, tidal-flat carbonates, however, extended well beyond the Appalachian area and occupied more than 975,000 km2 throughout eastern and central Laurentia (Figure 15); so extensive were these carbonates that they have been informally referred to as ‘‘The Great American Bank’’ (Demicco and Mitchell, 1982). Overall, Cambrian parts of the Appalachian section represent regional transgression toward the continental interior that is manifest in three, broad facies belts that migrated cratonward in time (Palmer, 1971).

Figure 15 Distribution of the Cambro-Ordovician ‘‘Great American Bank’’ in eastern Canada and eastern and central United States (brick pattern) relative to the present-day Appalachian Basin (dark outline; see Figure 14). The carbonate bank was not an Appalachian Basin phenomenon, but rather a paleogeographic and paleoclimatic phenomenon related to passive-margin status. Most of the carbonates are reef-related limestones near the eastern margin of the bank and grade westwardly into peritidal dolostones. Parts of the bank may occur below the Blue Ridge and Piedmont megathrust, east of the present basin margin. H, approximate Humber margin.

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Lower siliciclastic facies are called the Inner Detrital Belt and represent transgressive reworking of residuum on the underlying Precambrian surface in nearshore, terrestrial settings and in high-energy, littoral environments. The succeeding Middle Carbonate Belt reflects reduced clastic influx in slightly deeper, shallow-shelf, platformmargin to platform-lagoonal environments. Where preserved, the carbonate belt is succeeded by an Outer Detrital Belt that represents deep-water, continental-rise and basin sedimentation; it is characterized by dark graptolitic shales, interbedded with thin, dark, micro-grained limestones, conglomerates, breccias, turbidites, and siltstones, commonly with volcaniclastic input (Palmer, 1971; Reinhardt, 1977). Outer Detrital Belt sediments are largely preserved in northern parts of the area in Vermont, New York, and Massachusetts, but are perhaps best known from the Frederick, Kinzers, and Conestoga formations of Maryland and Pennsylvania. This pericratonic, passive-margin sequence persisted throughout the Appalachian area and into central parts of the Laurentian craton until the Middle–Late Ordovician transition. The units are widespread and relatively uniform across large distances and reflect the overall stability of the rifted, Laurentian craton margin (Figure 14) and interior. The eastward thickening of the sequence no doubt represents gentle thermal subsidence along the margin, whereas other facies variations are largely attributable to eustatic variation and local subsidence on rift-related structures. A similar passive-margin sequence and facies belts developed in New England and the adjacent Canadian Maritime provinces in the Quebec reentrant and on western Newfoundland parts of the St. Lawrence promontory on what has been called the Humber margin (van Staal, 2005). Basal, Neoproterozoic to Middle Cambrian, platform and slope, clastic sequences include the Oak Hill and Saint-Roch groups in Quebec and the Curling and Labrador groups in Newfoundland; Middle Cambrian to Lower Ordovician, platform and slope, carbonate-containing units are included in the Beekmantown, Philipsburg, and Trois Pistoles groups in Quebec and in the Port au Port, St. George, Cowhead, and Northern Head groups in Newfoundland, although many of the rocks contained in these groups are not so easily associated as they now occur in allochthonous nappes and thrust sheets (see Lavoie et al., 2003, and Lavoie, Chapter 3). Interestingly, by Early Cambrian time, a narrow seaway called the Notre Dame trough or Humber seaway (Figure 3) had developed between the Humber margin and the outboard Dashwoods microcontinent (Waldron and van Staal, 2001; van Staal, 2005), and by Middle Cambrian time, east-directed subduction had begun east of the Dashwoods microcontinent (van Staal, 2005), deforming the microcontinent, producing an infant arc, and flooding deeper parts of the Humber seaway with clastic sediments that began lapping westward onto carbonate-slope sediments from the Humber margin (James et al., 1989). At least in this region, early convergence events associated with the closure of Iapetus had begun by Middle Cambrian time and overlapped temporally with Humber passive-margin development.

9. Two Orogenic Cycles and the Origin of the Appalachian Foreland Basin The Appalachian Basin is a composite foreland basin that formed during the ocean-basin-closure phase of the post-Grenville Wilson cycle (Wilson, 1966; Dewey and Burke, 1974). Although the Neoproterozoic to Early Cambrian synrift and Early Cambrian to Middle Ordovician postrift, passive-margin deposition described above reflects a part of the same post-Grenville, Iapetan (Appalachian) Wilson cycle, these sediments are practically not a part of the foreland-basin sequence. In fact, during the final, Early Ordovician, passive-margin phase, a homogeneous blanket of shallow-water carbonates, The Great American Bank mentioned previously, stretched from the pericratonic, Appalachian continental margin into the central Laurentian craton (e.g., Cook and Bally, 1975; Neuman, 1976) (Figure 14). During the Early–Middle Ordovician transition, however, the nearly continent-wide Owl Creek unconformity developed (Figures 4 and 14), and the Appalachian continental shelf and adjacent parts of the foreland experienced profound sedimentologic and stratigraphic differentiation that signaled tectonic mobilization and initiation of the Appalachian foreland basin. As a result, the widespread and relatively uniform nature of pre-Owl Creek units (Figures 14 and 15) abruptly gave way to a complex, post-Owl Creek unit mosaic that reflects major tectonic influence and characterizes Appalachian Basin stratigraphy for the rest of Paleozoic time (Figure 14). The nearly continuous tectonic mobilization of the Appalachian margin during the rest of Paleozoic time occurred in two, phased orogenic mega-cycles (e.g., Ziegler, 1989). In the Appalachian Basin, these cycles are apparent as generalized, second-order, sea-level cycles (Figure 4, Caledonian and Variscan–Hercynian). The earlier, Ordovician–Silurian, Caledonian orogenic cycle on the Appalachian margin of Laurentia represents closure of the Iapetus (Theic) Ocean through subduction and collision involving island-arc complexes and peri-Gondwanan terranes during formation of the new continent, Laurussia; it was accomplished during two orogenies — the earlier Taconian, progressing from south to north in the south and central Appalachian areas, and the later Salinic, progressing from north to south. The following, Variscan– Hercynian orogenic cycle, in contrast, reflects closure of the Rheic Ocean during collision with Gondwana and

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the related transpressional adjustment of peri-Gondwanan terranes to form Pangea; it occurred on the Appalachian margin of Laurussia during two orogenies — the earlier Acadian (including the Neo-Acadian), progressing from north to south, and the later Alleghanian, probably progressing from south to north in its later phases. Before moving on, use of the words, ‘‘Taconian’’ and ‘‘Taconic,’’ needs to be clarified to avoid confusion, as the words have been used differently by various workers. The word Taconian is used in the sense of Kay (1969) to indicate all Ordovician–earliest Silurian deformation along the Appalachian margin regardless of timing and kinematic style. In contrast, the terms ‘‘Blountian’’ and ‘‘Taconic’’ are used to represent, respectively, Middle– Late and Late Ordovician tectophases within the Taconian orogeny (Kay and Colbert, 1965), although some have effectively elevated these two Taconian tectophases to the rank of orogenies (e.g., Drake et al., 1989) while others have used the term ‘‘Taconic’’ interchangeably with the more encompassing term ‘‘Taconian’’ (e.g., Hatcher, 2005) used herein. Comparative information on the progression of these orogenic cycles in the New England and Canadian Atlantic provinces of the northern Appalachian region are presented in Chapter 3 (Lavoie, Chapter 3).

9.1. Middle Ordovician–Early Devonian Caledonian orogenic cycle (Tippecanoe sequence, 472–411 Ma) 9.1.1. Taconian orogeny and foreland-basin sedimentation (Middle Ordovician–Early Silurian, 472–B436 Ma) The Taconian orogeny is usually thought to have begun near the Early–Middle Ordovician transition, and its initiation is marked nearly everywhere in the Appalachian region by a major unconformity. However, in many places throughout the Appalachian region, evidence exists for a latest Cambrian–earliest Ordovician, pre-Arenig to Arenig deformational and thermal event, and locally, an even earlier unconformity. These phases of tectonism have been called the Penobscottian disturbance, Penobscottian orogeny, or Penobscot phase of the Taconian orogeny (Neuman, 1967; Rodgers, 1967, 1970; Osberg, 1983; Hatcher, 1987, 1999; Skehan, 1988; Hall and Roberts, 1988, Rast, 1989). According to Hatcher (1987, 1999), the Penobscot phase most likely represents initiation of eastward subduction of oceanic parts of Laurentia below island-arc complexes that formed adjacent to the margin, but Hibbard and Samson (1995), van Staal et al. (1998), and van Staal (2005) have countered that the Penobscot phase most likely reflects a convergence event at the leading edge of Gondwana on the opposite side of the Iapetus Ocean. The event was unrelated to Laurentia except that Penobscot-influenced terranes were later exported from Gondwana (see Veevers, 2005) and accreted to Laurentia, and the ‘‘exported’’ nature of the deformation was not initially appreciated; the coeval Grampian orogeny is recorded in Laurentian parts of Scotland and Ireland (Rast and Crimes, 1969) and may reflect similar events. The initial convergence at the eastern margin of Laurentia, which marked the inception of the Taconian orogeny, had a rather complex history and paleogeography, perhaps not unlike that seen in the southwestern Pacific today (van Staal et al., 1998) with different styles of convergence going on together, or at different times and places. On the Laurentian margin, earliest phases of convergence had already begun in Cambrian and Early Ordovician time far east of the actively developing passive margin based on possible evidence in the southern Appalachians (McClellan et al., 2005a, b) and certain evidence in the northern Appalachians (van Staal, 2005) where the tectonism was already influencing passive-margin sedimentation (Lavoie et al., 2003; Lavoie, Chapter 3). By the Early–Middle Ordovician transition, however, collision of the southern and central Laurentian margin with an offshore arc had begun, and large parts of the passive, Appalachian continental shelf and slope were uplifted and eroded. Similar erosion also extended across most of cratonic Laurentia due to continental braking or bulge uplift and migration, resulting in a major unconformity variously called the St. George, postSauk, Owl Creek, sub-Tippecanoe, or Knox unconformity. In the Appalachian area, this unconformity reflects the beginning of the Taconian orogeny, the transition from a passive to a convergent margin, and inception of the Appalachian foreland basin; it also marks the first major tectonic differentiation of eastern Laurentia from a broad platform and ramp (Figure 15) to a series of basins and uplifts since its Late Proterozoic formation, as well as the related sedimentary differentiation of eastern and central Laurentia from the widespread, uniform, carbonate facies of The Great American Bank (Figure 15) to a variety of deep-to-shallow and clastic-to-carbonate facies (Figures 7–10). The unconformity is commonly characterized by karst, including sub-unconformity caves and sinkholes filled with breccias and may exhibit relief greater than 140 m; locally the surface was mineralized (Hoagland, 1971; Mussman and Read, 1986). The occurrence, timing, and locally angular nature of the unconformity clearly point to a major tectonic component, but Mussman and Read (1986) also argued for a eustatic component, and the bathtub-ring pattern of conformable strata from the deepest part of the new basin in central and eastern Pennsylvania (Figure 16) supports this interpretation. Although the unconformity is widespread across Laurentia, coeval unconformities are present on parts of Siberia, Baltica, and southeast Asia,

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Figure 16 Distribution of the post-Sauk, Owl Creek, sub-Tippecanoe, or Knox unconformity (see Figures 4 and 14) in the present-day Appalachian Basin (darkened area). Note the areas behind each promontory (dark arrows) where no unconformity developed and the large embayment in unconformity distribution, or partial ‘‘bathtub-ring’’ pattern that suggests eustatic in£uence.

suggesting to Hatcher (2007) that the unconformity and accompanying orogenies reflect a global pattern of crustal and upper mantle reorganization. The earliest deposition on the surface occurred in western parts of the basin as the St. Peter Sandstone, an eolian sand reworked into well-sorted, strandline deposits by newly transgressing seas. The St. Peter is sporadic in occurrence and is no more than 15-m thick; it is restricted to parts of the Rome trough in eastern Kentucky and West Virginia, suggesting accumulation in a structural low (Cable and Beardsley, 1984; Shaver, 1985; Humphreys and Watson, 1997). Although the collision zone extended along the entire eastern margin of Laurentia (e.g., Osberg, 1983; Faill, 1997a), convergence was diachronous and heterogeneous in kinematic style (e.g., Rodgers, 1970, 1971; Hiscott, 1984; Drake et al., 1989; Bock et al., 1996; Mac Niocaill et al., 1997). In the northern Appalachian area, convergence with outboard, rifted-margin elements or the margin itself seems to have been nearly continuous from Middle Cambrian to Early Silurian time (van Staal et al., 1998; van Staal, 2005), whereas in the central and southern Appalachians, there was a progressive, northeastward shift of convergence in time

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(e.g., Kay and Colbert, 1965; Shanmugam and Lash, 1982; Bradley, 1989; Pollock, 1989; Ettensohn, 1991), and irregularities or promontories along the margin localized deformation at various times and places (Dewey and Burke, 1974; Dewey and Kidd, 1974), generating three Taconian tectophases (Figure 4, cycles 1–3) with respective foreland-basin deposits in the larger Appalachian Basin (Ettensohn, 1991, 1994, 2004; Ettensohn and Brett, 2002). Lavoie (1994) has noted similar Taconian tectonic diachronism between the Quebec reentrant and St. Lawrence promontory. The earliest phase of Taconian orogeny (Figure 4, cycle 1) in the central and southern Appalachians was concentrated along southwestern parts of the margin near the Virginia and Alabama promontories (Figures 7), during Late Middle to Early Late Ordovician (Late Llanvirn–Mid-Caradoc; Late Whiterockian–Late Mohawkian) time and reflects collision of the margin with an island arc represented by the Dahlonega terrane and related volcanics from the Hillabee Greenstone (McClellan et al., 2005a, b), and/or collision with various crustal fragments or microcontinents at one or both promontories (Vick et al., 1987; Faill, 1997a). The ages of felsic extrusives and intrusives from the Dahlonega terrane and Hillabee Greenstone, and from several other possible nearby plutonic bodies, are all of Middle Ordovician age and are subduction-related (McClellan et al., 2005a), providing evidence for an arc that had previously gone ‘‘missing’’ for this phase of the Taconian orogeny (e.g., Drake et al., 1989; Windley, 1995). This early phase of Taconian orogeny is commonly called the Blountian tectophase and is represented in southern parts of the Appalachian Basin by a complete tectophase cycle (Figures 8 and 9); locally, bentonites are present near the base of the cycle (Adams et al., 1926). The cycle begins with basal transgressive carbonates and overlying, dark, graptolitic shales (Athens, Rockmart, Columbiana, Blockhouse, Paperville, Liberty Hall formations; 5–450 m thick) and is overlain by a westwardly prograding, relaxational clastic wedge with redbeds, sometimes called the Blount delta (e.g., Dunbar, 1960; Rogers, 1961a, b; Chowns and McKinney, 1980) (Figures 8 and 9). Although much of this clastic wedge was foreshortened or destroyed by later tectonism, the wedge probably extended no more than 125 km from its source area and attained a thickness of at least 2,700 m (Colbert and Kay, 1965; Cook and Bally, 1975; Chowns and McKinney, 1980; Patchen et al., 1985a). The Sevier foreland basin (Figures 8 and 9) is the earliest expression of the Appalachian foreland basin and records the preserved distribution of basinal dark shales and overlying clastic sediments. However because the Appalachian area was still located in the evaporative, subtropical, trade-wind belt, shallow-water, Blackriverian carbonates (Little Oak, Chambersburg, Stones River, parts of the Chickamauga, Lowville, High Bridge, Black River groups), 50–450 m thick, predominate in parts of the larger Appalachian Basin area (Figure 10) not affected by major subsidence or clastic influx. By Mid-Caradoc time, convergence had progressed farther northeastward along the margin such that the next tectophase (Figure 4, cycle 2) was mainly localized at and near the New York promontory (Figure 7). However, peak metamorphic dates from the eastern Blue Ridge of about 460 Ma (Busch et al., 2002; Moecher et al., 2004), also Mid-Caradoc in age, indicate that convergence was also occurring farther to the south near the Virginia promontory, perhaps associated with initial docking of the Carolina terrane (Vick et al., 1987; Noel et al., 1988; Hibbard, 2000; Dennis, 2006, 2007) (Figure 7). This Late Ordovician (Mid-Caradoc–Ashgill; Late Mohawkian– Cincinnatian) phase of orogeny involved subduction and eventual collision with island-arc terranes and is known as the Taconic tectophase; its initiation is represented by an unconformity at the Blackriverian–Chatfieldian (Rocklandian) boundary, which is best developed in areas behind the New York promontory (Ettensohn, 1994) (Figures 8 and 10). The boundary and inception of the tectophase are also marked by a series of bentonites, present in the basin from Alabama to New York (Adams et al., 1926; Haynes, 1994; Kolata et al., 1996; Mitchell et al., 2004). This tectophase resulted in a widening of the foreland basin that is indicated by the distribution of the Martinsburg, Reedsville, Utica, and Antes dark shales (Figure 7), with thicknesses up to 2,400 m. Some of this increased widening may be related to a Mid-Chatfieldian (Mid-Caradoc) east-to-west polarity shift in Taconic subduction (Coakley and Gurnis, 1995; Karabinos et al., 1998; Karabinos and Hepburn, 2001) that would have generated more regional compression and greater concomitant loading. Consequently, the Martinsburg foreland basin seems to have expanded at this time, in part reflecting subsidence along reactivated basement structures (Ettensohn et al., 2004a). Moreover, more proximal and coarser Martinsburg and Reedsville clastic sediments, like those occupying similar cycle positions in the Sevier and Knobs formations during the Blountian tectophase, not only exhibit prominent northwestward, perpendicular-to-strike paleocurrent structures, but also prominent, parallel-to-strike, north and south, paleocurrent directions (Meckel, 1970; Shanmugam and Walker, 1978, 1980; McBride, 1962; McIver, 1970) that suggest partition of the proximal foreland basin into local sub-basins by reactivated basement structures. The basin was subsequently filled with one of the most extensive, relaxational, Paleozoic clastic wedges, the upper parts of which contain extensive, marginal-marine redbeds (Queenston, Juniata, and Sequatchie formations) that have been called the Queenston Delta (Figures 8–10, 14). Because upper parts of the Queenston wedge are eroded, it must have been substantially thicker than the 380 m that are still preserved in places

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(Drake and Epstein, 1967; Colton, 1970; Patchen et al., 1985b). Although thickest parts of the wedge are centered on the New York promontory, the wedge extended at least 650 km cratonward into the Michigan Basin and nearly 1,700 km along strike from northeastern Tennessee, where it overlapped the Blount wedge, into southern Canada (Kay and Colbert, 1965; Poole et al., 1968; Colton, 1970; Sanford, 1993b; see Lavoie, Chapter 3) (Figures 8 and 17). In the most distal, western and southern parts of the basin, clastic-wedge facies grade into shallow-water, argillaceous carbonates and calcareous mudstones (Weir et al., 1984; Chowns and McKinney, 1980; Neathery and Drahovzal, 1985). The Queenston wedge is truncated by a widespread unconformity at the Ordovician–Silurian transition, called the Cherokee Discontinuity by Dennison and Head (1975) (Figures 4 and 14). The unusually widespread distribution of underlying redbeds (Dennison, 1976), the worldwide nature of the unconformity (McKerrow, 1979), and its coincidence with glacial deposits in Gondwana (Dennison, 1976; Hambrey, 1985; Caputo and Crowell, 1985; Bjorlykke, 1985), suggest that the unconformity was largely glacio-eustatic in origin. However, the facts that the unconformity becomes increasingly angular in northeastern parts of the basin and that there is an overlying tectophase cycle (Figure 10) strongly suggest tectonic influence (Rodgers, 1971; Liebling and Seherp, 1982; Ettensohn and Brett, 2002). In fact, evidence for Early Silurian intrusive activity, bentonites, and deformation near the St. Lawrence Promontory (e.g., van Staal, 1994; van Staal and de Roo, 1995; West et al., 1995; Bergstro¨m et al., 1997) and farther south (Hibbard, 2000; Hibbard et al., 2002), as well as flexural modeling of large stratigraphic sequences (Quinlan and Beaumont, 1984; Tankard, 1986; Ettensohn and Brett, 2002), all indicate the presence of a final tectophase of Taconian orogeny (Figure 4, cycle 3) concentrated at the

Figure 17 Major Caledonian-cycle clastic wedges. Distribution of major Taconian and Salinic clastic wedges on southeastern Laurentia. Arrowheads superimposed on Taconian map represent major paleocurrent directions as indicated in Meckel (1970), Shanmugam and Walker (1978, 1980), McBride (1962), and McIver (1970); SL, St. Lawrence promontory; NY, NewYork promontory; V,Virginia promontory; AL, Alabama promontory; A, Appalachian Basin; BW, Black Warrior Basin; I, Illinois Basin; M, Michigan Basin; C, Cincinnati arch (from Ettensohn, 2004).

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St. Lawrence Promontory, and possibly at the Virginia Promontory, in latest Ordovician and Early Silurian time. In the area of the Quebec reentrant and St. Lawrence promontory, this phase of orogeny is largely related to the accretion of Ganderia, a peri-Gondwanan microcontinent, and intervening arcs to Laurentia (van Staal and de Roo, 1995; Wilson et al., 2004; van Staal, 2005; see Lavoie Chapter 3). A Lower Silurian (lower Llandovery; lower Alexandrian) clastic wedge, preserved as the Medina Group (Meckel, 1970; Brett et al., 1990; Goodman and Brett, 1994), was generated, but only distal parts of the wedge are apparently preserved in the northern Appalachian Basin (Figure 10), where they attain a maximum thickness of 25 m. In north and central parts of the basin, basal transgressive parts of the sequence are typically represented by marginal-marine to shoreface, orthoquartzitic sandstones (Tuscarora, Clinch, and Whirlpool sandstones, typically 10–180 m thick), which are overlain by dark, basinal shales of the Power Glen and Cabot Head formations (0–25 m thick), whose distribution defines the maximum extent of the foreland basin (Figures 7 and 14); overlying, relaxational, clastic units are typically eroded and poorly preserved. The basal orthoquartzitic sandstones may reflect coastal, marine reworking of an eolian regolith in seasonally arid conditions (Cecil et al., 2004). In southern parts of the Appalachian Basin during Early Silurian time, a smaller and separate sub-basin developed with generally deeper water environments than those characterizing coeval, orthoquartzitic sandstones to the north. In eastern Tennessee, the dark shales, siltstones, and sandstones of the Rockwood Formation filled the basin, whereas in Georgia and Alabama, iron-rich shales and sandstones of the Red Mountain Formation predominate (Chowns and McKinney, 1980). The apparently separate nature of this sub-basin and its sedimentary infill may reflect deformational loading and uplift of source areas that accompanied continued transpressional convergence of the Carolina terrane with the Laurentian margin near the Virginia Promontory (Hibbard, 2000; Hibbard et al., 2002; Dennis, 2007) (Figures 7 and 18). Both exotic Carolina and Gander terranes (Figures 7 and 18), are thought to have been peri-Gondwanan microcontinents (e.g., Hibbard and Samson, 1995; Dennis, 1995; Hibbard, 2000; van Staal, 2005; Hatcher, 2005) that calved off the margin of Gondwana in a Late Cambrian–Early Ordovician episode of edge-driven terrane export (Veevers, 2005). Export of these terranes toward Laurentia left behind the widening Rheic Ocean and is

Figure 18 Generalized, Silurian tectonic setting of the Appalachian area, showing convergence of the ‘‘Reading Prong’’ (Avalonia) with the northern Laurentian margin to generate the Salinic disturbance. Note the southward migration of Salinic foreland basins during two tectophases based on the distribution of dark-shale basins (see Figures 14 and 19). The Carolina terrane (C) had apparently converged with the Laurentian margin in Ordovician time, but by Silurian time had begun to disperse northward, leaving possible successor basins to the west of it, which may have persisted into Early Devonian time. The area of open sea between Avalonia and the Taconian margin of Laurentia may also be related to the Merrimack trough and later successor-basin development. G, Gander terrane; V,Virginia promontory; NY, NewYork promontory; SL, St. Lawrence promontory. Adapted from Ettensohn and Brett (1998).

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almost certainly related to the nearly coeval onset of subduction below the Laurentian margin in the Iapetus Ocean (Murphy et al., 2006). Both terranes show evidence of earlier, peri-Gondwanan tectonic activity. On the Carolina terrane, the latest Proterozoic–Early Cambrian (610–520 Ma) Virgilina orogeny has been recognized, and on the Gander terrane, the previously mentioned Cambrian–Early Ordovician Penobscot orogeny has been identified. Moreover, in the Piedmont terrane, southeast of Carolina, the latest Proterozoic–Early Ordovician (550–480 Ma) Potomac orogeny has been recognized locally. Although large parts of the Piedmont terrane are now being separated into distinct peri-Laurentian subunits (e.g., Hatcher, 2005; Dennis, 2006a, b, 2007), parts of the terrane may still include old peri-Gondwanan arcs that accreted earlier to Carolina. Although it was thought by some that these orogenies had occurred on or very near Laurentia, recent evidence suggests that these events occurred elsewhere and are unrelated to the Laurentian–Iapetan realm (Hibbard and Samson, 1995). 9.1.2. Salinic disturbance and foreland-basin sedimentation (Early–Late Silurian, B436–417 Ma) The term ‘‘Salinic’’ or ‘‘Salinian’’ was first introduced by Boucot (1962) to designate a prominent, regional, Early–Late Silurian unconformity in New England; at the time, it had no orogenic or deformational implications. Subsequently, the unconformity was found to extend into eastern Canada, and by 1970, had been related to minor uplift or tectonic disturbance based on the presence of an angular unconformity and a clastic wedge (red clastics in the Salina Group of New York, from which it apparently takes its name; Rodgers, 1970). Rodgers (1970) even went on to indicate that the Salinic was probably a lesser counterpart of the Caledonian orogeny. Later, the Salinic came to represent a distinctive, pre-Acadian, tectono-metamorphic event (e.g., Waldron et al., 1998; Tremblay and Castonguay, 2002; Wilson et al., 2004; Lavoie and Asselin, 2004; see Lavoie, Chapter 3) with tectonic and sedimentary implications as far south as the central Appalachians (e.g., Ettensohn, 1994; Ettensohn and Brett, 1998, 2002; Dennis, 2006a, b, 2007). Unconformity development, however, also coincided with a period of Middle and Late Silurian sea-level decline (Wilson et al., 2004; Lavoie and Asselin, 2004; see Lavoie, Chapter 3). The final, Early Silurian, Taconian tectophase cycle (Figure 4, cycle 3) was almost immediately followed by an apparently unrelated Early Silurian (latest Llandovery–Early Wenlock; Clinton) cycle in northern parts of the Appalachian Basin. The cycle reflects the first part of a two-phase (Figure 4, cycles 4 and 5), orogenic event, called the Salinic disturbance (Boucot, 1962) (Figures 4, 18, and 19), which consisted of at least two, southwestwardly migrating, Early–Late Silurian (Late Llandovery–Pridoli; Niagaran–Cayugan) tectophases (Ettensohn and Brett, 1998), evidence for which is best preserved in the northern Appalachian Basin (Figure 18). Unlike Taconian tectophases, the Salinic cycles are more carbonate-rich and migrated southward in time and space (Ettensohn and Brett, 1998, 2002) (Figures 18 and 19). The disturbance apparently reflected southward extension of Caledonian tectonism during the convergence of Avalonian terranes in the ‘‘Reading Prong’’ of Baltica with the northern Appalachian margin of Laurentia (Mckerrow and Ziegler, 1972; Ettensohn, 1994; Ettensohn and Brett, 1998; Wilson et al., 2004) (Figure 17). Van Staal (2005) has suggested that the orogeny represents the final collision of

Figure 19 Schematic southwest-northeast section partially parallel to basin strike, showing two tectophase sequences generated in the Appalachian Basin during the Salinic disturbance. Note the southwestward migration in time of black-shale basins that re£ects a similar trend in Salinic convergence (Figure 18). The Williamson-Rose Hill tectophase sequence did not go to completion and contains no subaerial components, but the second, Rochester-Bloomsburg-Salina sequence went to completion and ends with Bloomsburg-Vernon redbeds and Salina evaporites (see Figure 14). No vertical scale intended (from Ettensohn, 1994).

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Ganderia with Laurentia, but Salinic deformation is so pervasive across Ganderia as to suggest that Ganderia might have been caught in a pincers movement between Laurentia and Avalonia (Figure 18). Avalonia played an important role in Appalachian tectonics and developed as a series of microcontinents that formed as an ensialic, volcanic arc on the Neoproterozoic, west African margin of Gondwana and subsequently rifted from it in Late Cambrian to Early Ordovician time during closure of the Iapetus sea (Secor et al., 1989; Nance and Murphy, 1994; Nance et al., 2004; Murphy et al., 2006; Thompson et al., 2007); with Baltica, it subsequently converged on the northern Laurentian margin to generate the Caledonian (Scandian) orogeny in the north (see Roberts, 2003) and the Salinic disturbance to the south. The accretion of Baltica to Laurentia generated a new, combined continent called Laurussia (see Ziegler, 1989). Although Rodgers (1987) concluded that the disturbance was concentrated near the St. Lawrence Promontory, tectophase cycles from the Appalachian Basin indicate that the effects of convergence were prominent as far south as the Virginia Promontory. Development of a regional angular unconformity in New York, central Pennsylvanian, and southern Ontario parts of the basin and an equivalent disconformity on adjacent parts of the craton (Ettensohn, 1994) marks inception of the first tectophase, and dark shales in the Williamson and equivalent parts of the Rose Hill formations define the extent of the foreland basin (Ettensohn and Brett, 1998) (Figures 14, 18, and 19). In other parts of the Appalachian area, fine-grained marine clastic sediments in the Rose Hill, Clinton, and Crab Orchard formations, ranging from 10 to 290 m in thickness, reflect an equivalent transgressive interval. No subaerial facies were preserved in the cycle, which apparently did not go to completion (Figure 19). From New York to Alabama, Lower Silurian rocks in the Appalachian Basin, whether reflecting latest Taconian or Early Salinic tectonism, are known for their redbeds and iron-rich shales, sandstones, and carbonates, some of which have served as iron ores. These iron-rich sediments are thought to reflect the deep subaerial weathering of Taconian highlands, coincident with development of the Cherokee discontinuity in semiarid to seasonally wet climates (Retallack, 1993; Cecil et al., 2004) (Figure 4). Resulting iron-rich regolith and dissolved iron were subsequently transported into the warm, shallow, coastal seas, where the red sediments were deposited, and iron in solution was precipitated, some of it cementing clastic components and replacing shallow-marine carbonates (Folk, 1960; Ziegler and McKerrow, 1975; Chowns and McKinney, 1980). The second Salinic tectophase cycle, however, is complete and is well-known for its Bloomsburg redbeds and related Salina evaporites (Ettensohn, 2004) (Figure 19). The tectophase cycle is defined at the base by an Early Wenlockian (Mid-Clinton) unconformity overlain by transgressive, shoreface deposits of the Keefer Sandstone and equivalents, as well as by the overlying, dark, basinal Rochester Shale (Figure 19). The succeeding Bloomsburg redbeds form the relaxational clastic wedge and attain a maximum thickness of 550 m; they extend at least 600 km along strike and about 300 km cratonward (Colton, 1970; Cook and Bally, 1975; Patchen et al., 1985b) (Figure 19). Although upper parts of the Shawangunk Formation in southeastern New York, northern New Jersey, and northeastern Pennsylvania may represent very proximal equivalents of both Salinic clastic wedges (Figure 19), preserved parts of the wedges largely reflect the distal, fine-grained parts of delta complexes with major source areas to the northeast in present-day New England (Ettensohn, 1994, 2004) (Figure 19). Late Silurian (Late Ludlow–Pridoli; Late Lockportian–Cayugan) time saw the only major episode of Appalachian evaporite deposition, which is largely restricted to northern and north-central parts of the basin (Figures 19 and 20). The evaporites occur as salt and anhydrite associated with shaly dolostones in parts of the Salina Group, with thicknesses from 120 to 800 m thick (Patchen et al., 1985b). The evaporites appear to occur in basins largely defined by tectonic elements, like the Cincinnati-Kankakee-Algonquin arch system and a likely Salinic forebulge, which hosted shallow-water, carbonate ramps, and at various times, ‘‘reefal’’ or biohermal banks that apparently restricted basin waters (Kay and Colbert, 1965; Brett et al., 1990; Goodman and Brett, 1994) (Figure 20). Evaporitic rocks in the basin, and elsewhere throughout the world, are characterized by fourth- and fifth-order, shallowing-upward cycles (e.g., Johnson, 1996; Cecil et al., 2004) that probably reflect Milankovitch cyclicity. Overall, however, tectonic elements, combined with presence in the evaporative tradewind belt (Scotese, 2003) and a likely rain shadow behind Taconian and Salinic mountains (Grabau, 1913), were important factors in Silurian deposition of carbonates and evaporites in the northern Appalachian Basin. In the central and southern parts of the basin, Upper Silurian rocks were commonly destroyed by Devonian erosion, but where preserved, dolomitic carbonates predominate. 9.1.3. Helderberg transition interval and related basin sedimentation (latest Silurian–Early Devonian, 417–411 Ma) Conformably to unconformably overlying Bloomsburg-Salina rocks in the central and northern Appalachian Basin is a mixed-carbonate-clastic sequence of latest Silurian–Early Devonian age (Pridoli–Lochkovian–Early Pragian; latest Cayugan–Early Ulsterian), called the Helderberg Group (Figures 4, cycles 6 and 7; and 14). Whether or not major sedimentary cycles in the group are of tectonic or eustatic origin and whether or not they

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Figure 20 Schematic, Late Silurian (Ludlow--Pridoli; Cayugan), paleogeographic map of Salinic evaporite basins, reef banks, and terrigenous, alluvial-plain sedimentation in the Appalachian foreland basin and adjacent areas. Development of basins and reef banks was probably controlled by bulge migration that reactivated regional basement structures, as well as by foreland subsidence and resulting basin yoking. Structures involved include the Algonquin arch (A), the Findlay arch (F), the Kankakee arch (K), the Cincinnati arch (C), the Iapetan Ohio-West Virginia hinge zone (O),Tristate block (T), and the Grenvillian Vanceburg-Ironton fault zone (V). Arrows point to downthrown or down-dipping sides. Reef banks along the Ohio-West Virginia hinge zone and Cincinnati-Findlay-Algonquin arches may re£ect prominent bulge positions. Bloomsburg-Vernon redbeds (B). Adapted from Kay and Colbert (1965).

are related to Salinic or Acadian tectonism or reflect a time of relative tectonic quiescence, are frequently debated questions. However, in regions far outboard of the foreland basin, where converging microcontinents had largely effected closure of the Iapetus Ocean, Late Silurian–Early Devonian time appears to have been a transitional period of extensional block faulting (Figure 21A), within-plate magmatism and regional uplift accompanying collapse of the Taconian orogen and development of a new Acadian transpressional tectonic regime. The extension that accompanied uplift resulted in a series of successor basins that received shelf and flyschoid sedimentation from Laurussian sources to the northwest and peri-Gondwanan sources to the south and southeast (Keppie and Dostal, 1994; Lavoie and Asselin, 2004; Wilson et al., 2004; Dennis, 2006a, b, 2007) (Figure 21A). In the northern Appalachians, deposition of Silurian–Devonian rocks occurred as overlap sequences in troughs now preserved as parts of three synclinoria (Gaspe´-Connecticut Valley, Merrimack-Aroostook, and Fredericton synclinoria) with evidence of rift-related sedimentation and igneous activity, as well as coeval structural features that support post-Salinic extension and successor-basin development (Osberg et al., 1989; Keppie and Dostal, 1994; Tremblay and Castonguay, 2002; Lavoie and Asselin, 2004; Wilson et al., 2004; Tremblay and Pinet, 2005; see Lavoie, Chapter 3). In New England, the rift-related depression formed an inland sea called the Merrimack trough (e.g., Osberg et al., 1989) (Figure 21A), which was largely infilled with dark shales, turbidites, and other deep-shelf clastics (Guidotti and Van Baalen, 2001; Lavoie and Asselin, 2004; Lavoie, Chapter 3 ), but farther to

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Figure 21 Generalized, schematic representations of the foreland and pro-foreland, tectonic settings during the (A) SiluroDevonian, Helderberg transition period, and during (B) the Devonian--Mississippian, Acadian fourth tectophase, or NeoAcadian orogeny. (A) Shows pro-foreland collapse of the Taconian-Salinic orogen due to delamination of downgoing slabs under Gander and Avalon (A) terranes; horizontally lined pattern re£ects areas accreted to Laurentia/Laurussia from Middle Cambrian--Late Silurian time; rift basins (stipple) developed on Gander and Laurussian crust during transition and in¢lled with dark shales and turbidites; major basins: CG, Connecticut Valley-Gaspe¤; CM, Central Maine-Merrimack; F, Fredericton; CS, Cat Square. Basinal area between Laurussian (Humber) margin and Avalon terranes has been called the Merrimack trough or sea. (M) During later convergence, basin sediments were tectonized or locally thrust westward. In the foreland, H, the preserved Helderberg foreland basin; the southern basin ‘‘tail’’ inVirginia,Tennessee, Georgia, and Alabama marks occurrence of the thin Frog Mountain,Wildcat Valley, and Rocky Gap sandstones; B, Big Mtn.; and Ma, Mandata black shales are small north-migrating shale basins that may re£ect earlier sinistral transpression between Carolina and Laurussian margin; C, Carolina terrane. (B) Shows Late Devonian--Early Mississippian, pro-foreland, pincers movement of Avalon terrane (A) and Meguma between Gondwana and Laurussia and resulting dextral transpressional movement of Avalonia and Carolina (C). The belt of thin, wavy lines represents a zone of intense transpressional deformation and possible tectonic channeling, especially involving former rift basin sediments (Figure 21A). In the foreland, the Catskill wedge (Ca; coarse stipple) re£ects Late Devonian transpressional movement from the NewYork to the Virginia promontories; the Sunbury (S; dark dotted line)-Riddlesburg (R; dashes and dots with lined pattern) dark shales and the Price-Pocono wedge (P; ¢ne stipple) re£ect transpressional movement fromVirginia promontory southward. Note the prominent eastward migration of the Sunbury-Riddlesburg basin toward the NewYork and Virginia promontories (see also Figure 24), where tectonic activity was concentrated. The Reguibat promontory of Africa would eventually converge on the NewYork promontory with subsequent clockwise rotational movement of Gondwana.

the north, there is also evidence for transgressive, coastal-plain, reef, and carbonate-platform sedimentation on the northwestern trough margin (Lavoie, 1992a, b; Lavoie and Asselin, 2004; Lavoie, Chapter 3). In a somewhat similar situation from the central Appalachian area, Dennis (2006a, b, 2007) has reported the formation of the Cat Square successor basin and related magmatism at the same time, while the Carolina terrane rifted from the Laurussian margin in an episode of post-accretionary terrane dispersal (Figure 21A). Van Staal and de Roo (1995), Cawood et al. (1995), and Wilson et al. (2004) support a model of lithospheric delamination below the Taconian/Salinic accretion zone that could explain this period of uplift, magmatism, and orogen collapse. The model is potentially important, because it not only means that other successor basins might have been present along the former margin in this transition period, but also because separation of the downgoing Avalonian slab during delamination may have initiated the new transpressive regime that characterized the succeeding Acadian orogeny (Wilson et al., 2004). The main, Helderberg, foreland-basin depocenter was in the north-central and central Appalachian Basin from southern New York though central Pennsylvania and into eastern West Virginia, where the group ranges up to 140 m thick; on the western margins of the basin in Ohio, West Virginia, and Kentucky (Figure 21A), the group is only present locally as thin, eroded remnants of dolomitic carbonates. To the south, in contrast, from western Virginia to Alabama, equivalents are present in thin, variably aged sandstone units like the Frog Mountain Sandstone of Alabama, the Wildcat Valley Sandstone of eastern Tennessee, and the Rocky Gap Sandstone of western Virginia (Figure 21A), all of which seem to represent episodic reworking on the fluctuating margin of a very shallow sea. In the main Appalachian depocenter, however, deepening clearly occurred toward the center of the basin, where up to 60 m of chert and cherty shale in the Shriver Formation were deposited. Cecil et al. (2004) related these cherts to an abundance of eolian, quartzose dust, deposited from winds in the arid, trade-wind belt (Figure 4).

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The predominance of carbonates in the group has led most workers to conclude that deposition occurred during a time of tectonic quiescence between the Taconian (or Salinic) and Acadian orogenies (e.g., Laporte, 1969; Dorobek and Read, 1986; Smosna, 1988); however, evidence for structural reactivation in the area (Goodmann, 1988; Linn et al., 1990), bentonites in the sequence (Ver Straeten, 2004), and synchronous Silurian–Devonian magmatism in the area (Horton et al., 1995; Samson and Secor, 2000; Hibbard et al., 2002) suggest that coeval tectonism probably generated upland sources for the small clastic wedges in the central Appalachian sequence. Helderberg facies are very complex and have been further complicated by superimposed, small-scale eustatic cycles of likely glacial origin (Saltzman, 2002), but two, third-order, transgressive-regressive cycles of about 2 and 4 Ma, respectively (Figure 4, cycles 6 and 7), centered on small, local, dark-shale basins in east-central parts of the larger Appalachian Basin (Figure 21A), have been identified (Smosna, 1988). By latest Early Devonian (Late Emsian; Late Ulsterian, Deerpark) time, Helderberg carbonate deposition began waning as a sheet of wellwashed sand inundated the basin in high-energy, shallow-marine, and littoral environments at the front of a new transgression. These sands, now represented by the Oriskany and Ridgely sandstones (Figure 12), reflect the beginning of the new Kaskaskia Sloss sequence, the new Acadian orogeny, and the final Variscan–Hercynian, orogenic cycle of the Appalachian Wilson cycle (Figure 4). Along the margins of the Appalachian Basin, the Helderberg–Oriskany contact is unconformable, but toward the basin center, contacts become gradational (Head, 1974; Ettensohn, 1994). The two dark-shale basins (Big Mtn. and Mandata shales; Figure 21A) with associated clastics and carbonates appear to be small-scale, foreland-basin, tectophase cycles, and although the tectonic origins are uncertain, they may reflect two early pulses of loading due to south-to-north sinistral transpression between the Carolina terrane and the Laurussian margin (Figure 21A). However, three anomalous thrust-bound, low-grade metamorphic, sequences in the central and southern Appalachian area may also be related to the tectonic setting shown in Figure 21A. At least two of the sequences, including parts of the Ocoee Supergroup of Georgia and Tennessee (Figure 13) and the Talladega Group of Georgia and Alabama, have been interpreted to contain Upper Silurian– Devonian sediments, although younger sediments may also occur (Unrug and Unrug, 1990; Tull and Groszos, 1990; Tull, 1998, 2002; Unrug et al., 2000); the third sequence, the Mineral Bluff Group of Georgia and North Carolina, is Ordovician or younger (Tull et al., 1993). Most of these sequences are floored with trailing-margin basement or outer-platform, Early Paleozoic carbonates, which are unconformably overlain by deeper water, turbiditic, olistostromal, and locally cherty units that are more characteristic of successor-basin deposits than of other types of rift basins (Tull and Groszos, 1990; Tull et al., 1993; Tull, 1998; Unrug et al., 2000). In fact, Tull and Groszos (1990), Tull et al. (1993), Tull (1998), and Unrug et al. (2000) have suggested that these sequences were deposited in successor basins related to transtensional and transpressional events during Early Acadian orogeny. However, the likely Late Silurian–Devonian age of the sediments is too old to be Acadian in this part of the Appalachians, but the cherty nature of some units may reflect the same climatic conditions that supported Helderberg chert deposition in the central Appalachian Basin, and probable, Salinic, extensional successor basins are already known from the northern Appalachians at the same time (Osberg et al., 1989; Tremblay and Castonguay, 2002; Wilson et al., 2004; Tremblay and Pinet, 2005; Dennis, 2007). Most likely, transtensional and transpressional events involving the Carolina terrane, which was already proximal to the Laurussian margin by Early Silurian time (Hibbard, 2000; Hibbard et al., 2002) (Figure 18), generated the basins and their infill (Figure 21A), and in some cases, infilling may have continued into later Devonian and Mississippian time (McClellan et al., 2005a, b; Dennis, 2006a, b, 2007). In fact, Hatcher et al. (2003), Bream et al. (2005), and Dennis (2006a, b, 2007) have suggested the presence of a small remnant ocean or successor basin between the Carolina terrane and Laurussia (Figures 18 and 21A) in which sediments could have accumulated; some of these sediments were apparently reworked into an accretionary prism in front of accreting and/or dispersing Carolina (Bream et al., 2005; Dennis, 2006a, b, 2007), whereas accreting sediments from more distal parts of the basin, or from other basins, may have been thrust relatively unaltered cratonward during later orogeny. Most likely, the final deformation and current, inboard positions of these basin sediments are probably related to Alleghanian deformation. Nonetheless, from the end of Early Devonian instability until the earliest Mississippian advent of the final Acadian (Neo-Acadian) tectophase at the Devonian–Mississippian transition, a period of about 48 Ma (Pragian–earliest Tournaisian), the southeastern margin of Laurussia apparently became a stable, shallow-shelf setting (Tull, 2002). Although Late Silurian–Devonian deformation, metamorphism, and magmatism in the southern Appalachians (Horton et al., 1995; Samson and Secor, 2000; Hibbard et al., 2002; Esawi, 2004), as well as presence of likely Helderberg-age successor-basin sequences (Tull and Groszos, 1990; Tull et al., 1993; Tull,1998; Unrug et al., 2000; Dennis, 2006a, b, 2007), support the presence of Helderberg tectonism, its effects on the Appalachian foreland basin were mild, for only two small tectophase cycles and their accompanying clastic wedges developed on the southeast-central margin of the Appalachian Basin. Shallow-water carbonates predominate in most other parts of the basin.

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Whether or not Helderberg deposition was related to Salinic or Acadian tectonism, or was the product of a separate event, is debatable. Sedimentary evidence from the Appalachian Basin suggests that during both Salinic and Acadian orogenies, convergence was from north to south, but the disposition of Helderberg dark-shale basins and clastic wedges (see Smosna, 1988) suggests a south-to-north convergence, which follows the sinistraltranspressive mode of accretion suggested by Hibbard (2000), Hibbard et al. (2002), and Dennis (2006a, b, 2007) for the Carolina terrane (Figure 21A). It is also likely that former Late Silurian–Devonian successor basins in the area and the small Helderberg clastic wedges in the Appalachian Basin merely reflect a transitional phase of uplift, magmatism and extension during collapse of the Taconian–Salinic orogen and dispersal of the Carolina terrane from Laurussia (Dennis, 2006a, b, 2007), just like similar events reported from the northern Appalachians (e.g., Keppie and Dostal, 1994; Lavoie and Asselin, 2004; Wilson et al., 2004). Although Ordovician–Silurian convergence of the Carolina terrane was discussed previously as a Taconian event because of timing, in reality its accretion may have been entirely unrelated to Taconian, Salinic, or Acadian events except for the temporal overlap. Because of this uncertainty, Helderberg sedimentation in the Appalachian Basin is treated here as a transitional response to the end of the Caledonian orogenic cycle and the beginning of the Variscan–Hercynian cycle; it is included at the end of the Caledonian cycle (Figure 4).

9.2. Early Devonian–Permian Variscan–Hercynian orogenic cycle (Kaskaskia and Absaroka sequences, 411–251 Ma) 9.2.1. Acadian (Neo-Acadian) orogeny and foreland-basin sedimentation (Early Devonian–earliest Pennsylvanian, 411–315 Ma) During most of the Acadian orogeny, the Appalachian Basin area was still located in the arid, subtropical tradewind belt (Scotese, 2003), although by latest Devonian–earliest Mississippian time, dry-subhumid to humid conditions briefly returned due to global cooling and glaciation on Gondwana (Caputo and Crowell, 1985; Veevers and Powell, 1987; Frakes et al., 1992) (Figure 4) and probably in the Appalachian Basin (Cecil et al., 2004). By latest Mississippian time, however, the basin area had begun moving into the humid, tropical, equatorial zone (Scotese, 2003), and the Carboniferous–Permian glaciation of Gondwana was initiated. Sea level, however, apparently reached its maximum in Late Devonian and Early Mississippian time and thereafter declined throughout the remainder of the Carboniferous and Permian periods (Vail et al., 1977; Johnson et al., 1985; Dennison, 1989) (Figure 4). The earlier Salinic disturbance had apparently been generated by the westward convergence and/or docking of various Avalonian terranes (Figure 18), which had apparently been juxtaposed near the Laurentian/Laurussian margin since Ordovician–Silurian time (Phillips et al., 2003; Murphy et al., 2004). Initial accretion of the periGondwanan Carolina and Avalon terranes to Laurentia appears to have had an oblique, sinistral aspect (e.g., Hibbard, 2000; Malo and Kirkwood, 1995; van Staal and de Roo, 1995; van Staal et al., 1998; van Staal, 2005), and van Staal and de Roo (1995) confirmed that the switch from dominantly sinistral to dextral translation after late Early Devonian time indicates that Acadian transpression is kinematically unrelated to any Silurian–earliest Devonian convergence. Nonetheless, it was only because of likely delamination and slab loss (van Staal and de Roo, 1995; Wilson et al., 2004; Tremblay and Pinet, 2005) during the Helderberg transition (Figure 21A) that dextral, oblique convergence between Avalonian terranes and Laurussia (e.g., Ferrill and Thomas, 1988; Malo and Kirkwood, 1995; van Staal and de Roo, 1995; Trupe et al., 2003), became possible, perhaps under the influence of Meguma and/or Gondwanan impact by Late Devonian (Famennian, Chautauquan) time (Ziegler, 1989) (Figure 21B). Hence, it is this prominent change in translation polarity that marks the inception of Acadian orogeny in late Early Devonian (Pragian–Emsian; Mid-Ulsterian, Deerparkian) time, and it was the dextral transpressional accretion of these terranes from the northeast to the southwest along the Laurussian margin that generated the similarly migrating clastic wedges in the Appalachian Basin for which the Acadian orogeny is so well-known (e.g., Barrell, 1913, 1914a, b; Friedman and Johnson, 1966; Dennison, 1985; Woodrow and Sevon, 1985; Faill, 1985; Ettensohn, 1985a, b, 1987) (Figures 21B and 22). The Acadian clastic wedges are parts of four, foreland-basin, tectophase cycles, some with pronounced subcycles (Figures 4, cycles 8–11; and 12) that developed and migrated southward in time and space (Ettensohn, 1985a, 1994, 1998) (Figure 23). Initiation of the orogeny is marked in the Appalachian foreland basin by the preOriskany or Wallbridge Discontinuity, which also defines the base of Sloss’s (1963) Kaskaskia sequence and the beginning of the Variscan–Hercynian orogenic cycle (Figures 4, 12, and 14). The earliest docking was focused near the St. Lawrence promontory and Quebec reentrant in northern New England and the Canadian Maritime Provinces in late Early Devonian time, and as a consequence in the New England area, the earlier generated Merrimack trough (Figure 21A) became filled with an accretionary wedge of volcanic, volcaniclastic, and flyschoid deposits; these deposits were subsequently deformed as the Avalonian, Narragansett-St. John terrane

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Figure 22 Distribution of Variscan--Hercynian, Acadian (Neo-Acadian) and Alleghanian clastic wedges on southeastern Laurussia. Arrowheads re£ect major paleocurrent directions as indicated in Meckel (1970) (from Ettensohn, 2004).

Figure 23 Distribution of cyclic, Late Devonian (Frasnian--Famennian), third-tectophase (see Figure 12) black-shale basins in the larger Appalachian Basin area. Note the southward migration of shale basins that must re£ect the coeval southward migration of Acadian orogeny. Adapted from Ettensohn (1985a).

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docked with Laurussia (see McKerrow and Ziegler, 1972; Rodgers, 1981; Bradley, 1983; Osberg et al., 1989) (Figure 21B). Although sporadically preserved, formerly clastic-rich deposits, like those preserved in the Carabassett, Seboomook, and Littleton formations of New England as quartzites, slates, shists, and gneiss (Skehan, 1985), may have been deposited as a clastic wedge in a foreland basin, but the nature of that foreland basin is now uncertain because of metamorphism, deformation, and partial destruction by later uplift and erosion. At the same time, however, in preserved parts of the Appalachian Basin to the south, convergence and related loading at the New York and Virginia promontories produced a relatively thin, incomplete tectophase cycle (Figure 4, cycle 8), beginning with the very shallow-marine to littoral Oriskany and Ridgeley sandstones, 0–100 m thick, and the equivalent Glenerie Limestone of southeastern New York, which deepened upward into dark shales and overlying coarser clastic sediments in the Esopus-through-Schoharie formations (5–180 m thick) in eastern Pennsylvania and eastern New York and into the dark Needmore Shale of southern Pennsylvania, Maryland, eastern West Virginia, and west-central Virginia (5–60 m thick) (Figures 12 and 14). Cherty Huntersville equivalents of the Needmore (Figure 14) to the southwest may reflect local continuation of conditions associated with chert deposition during Helderberg time (see Cecil et al., 2004). Equivalents to the south from western Virginia to Alabama are again concentrated in thin, variably aged sandstone units like the Frog Mountain, Wildcat Valley, and Rocky Gap sandstones, which apparently reflect continued, episodic reworking across a very shallow, fluctuating basin margin. The second tectophase (Figure 4, cycle 9) apparently represents Middle Devonian (Eifelian–Early Givetian; Erian) docking with the New York promontory. The corresponding tectophase cycle begins with the widespread pre-Onondaga unconformity, overlain in much of the basin by the Onondaga Limestone and its equivalents, which may be 5–60 m thick and deepen upward into the widespread, black Marcellus Shale (10–250 m thick) and overlying coarser clastic sediments of the Hamilton Group (10–1,000 m thick) (Figures 12, 14, and 24); probable equivalents in southwestern parts of the basin are again concentrated as parts of the Wildcat Valley and Frog Mountain sandstones. At the top of the Onondaga or in the base of the Marcellus and their equivalents,

Figure 24 Schematic southwest-northeast section across the central Appalachian Basin in Kentucky and West Virginia showing the approximate positions of cyclic, Middle Devonian to Lower Mississippian, black-shale basins and intervening clastic wedges formed during the second to fourth Acadian tectophases (see Figures 12 and 23). Note the cratonward migration of black-shale basins in time that re£ects a coeval migration of Acadian deformation loading in the same direction. Transgressive black shales are largely restricted to more proximal foreland-basin settings, whereas regressive black shales represent the distal edges of coarser clastic wedges into a deeper cratonic basin. The upper Olentangy Shale represents the time of yoking between the Appalachian and Illinois basins (see Figure 1). The Sunbury and overlying Borden and Price-Pocono clastic wedges represent fourth-tectophase or Neo-Acadian tectonism. In contrast to underlying black-shale basins, the Sunbury black-shale basin migrated more than 100 km to the east, re£ecting more intense and restricted fourth-tectophase or Neo-Acadian deformational loading. Adapted from Ettensohn et al. (1988a).

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the Tioga bentonite, a few centimeters to 5 m thick, reflects major second-tectophase volcanism with a probable source in the Piedmont (Dennison and Textoris, 1970; Dennison, 1989; Ver Straeten, 2004). In contrast, the third tectophase (Figure 4, cycle 10) reflects the southward migration of deformation and impending collision with the Virginia promontory in late Middle to Late Devonian (Late Givetian–Famennian; Late Erian–Chautauquan) time (Ettensohn, 1985a) (Figure 21B). The tectophase cycle begins with a local unconformity, which is overlain by the Tully Limestone (0–85 m thick) or black shales in a relationship that has been called the Taghanic onlap (Johnson, 1970). During the rest of the third tectophase, this onlapping of the western margin of the Appalachian Basin continues through five cycles of black-shale and coarser clastics (Figures 12, 14, 23, and 24), and by the time of the fourth, or Huron–Dunkirk cycle, cyclic black shales and coarser clastic sequences overlapped the bounding Cincinnati arch and spread westward into the Illinois and Michigan basins (Ettensohn and Geller, 1987, Ettensohn et al., 1988a; Ettensohn, 1998). Ettensohn (1985a, 1987, 1994, 1998) has demonstrated that successive black shale-coarser clastic cycles migrate not only progressively cratonward, but also parallel to basin strike (Figure 23), tracking the prominent oblique or strike-slip component of the Acadian orogeny; each of the five, third-tectophase subcycles, moreover, may reflect a major phase of thrust loading during the tectophase. The overlapping effects of the second and third tectophases near the New York and Virginia promontories produced the clastic wedge that Barrell (1912, 1913, 1914a, b) first called the Catskill Delta (Figures 11, 14, 21B, 22, and 24). Preserved parts of the Catskill Delta complex near the New York Promontory are more than 3,000 m in thickness; whereas those parts of the complex centered on the Virginia promontory attain thicknesses in excess of 2,400 m (Colton, 1970; Patchen et al., 1985b). Major parts of the complex extend nearly 800 km along parts of the basin behind the New York and Virginia promontories (Figures 22 and 24) and exhibit several major lobes and intervening embayments (Dennison and Head, 1975; Dennison, 1985; Woodrow, 1985; Boswell and Donaldson, 1988). Although commonly called a delta, the Catskill complex is largely represented by the Catskill and Hampshire formations and reflects an alluvial coastal plain, dominated mostly by fine muddy sediments (Walker, 1971; Woodrow et al., 1973; Sevon, 1985; Dennison, 1985); deeper water, marine equivalents from turbiditic to shallow, open-marine settings in the second tectophase are included in the Hamilton Group and Mahantango Formation, and from the third tectophase, in the ‘‘Chemung,’’ Lock Haven, Brallier, Scherr, and Foreknobs formations (Dennison, 1970, 1985; Dennison and Wheeler, 1975) (Figures 14 and 24). For the most part, the wedge extended no more than 400 km from source areas, but locally, tongues of subaqueous sediment extended distally another 300 km across the Cincinnati arch into eastern parts of the Illinois and Michigan basins (e.g., Ettensohn et al., 1988a; Matthews, 1993) (Figure 24). The famous Bedford-Berea Delta (Pepper et al., 1954) (Figures 22 and 24) is probably best considered to be a part of the Catskill complex, and apparently represents a period of latest Devonian, lowstand reworking of the complex followed by transgression (Pashin and Ettensohn, 1987, 1995). Part of the lowstand at this time was no doubt related to ongoing Gondwanan glaciation (Cecil, 1990; Frakes et al., 1992; Eyles, 1993), but there is also evidence for coeval glaciation in the Appalachian Basin. In fact, in extreme eastern parts of the Appalachian Basin in a belt nearly 400 km long, from eastern Pennsylvania across western Maryland and into north-central West Virginia, up to 10 m of polymictic diamictites, laminites, and mudstones with dropstones occur in lower parts of the Rockwell and Spechty Kopf formations. The diamictites and related deposits have been interpreted to represent debris-flow deposits into lakes (Sevon and Berg, 1986; Sevon et al., 1997; Berg, 1999) or tillites from Acadian alpine glaciation (Brezinski, 1989; Cecil et al., 2002, 2004; Dennis, 2005a, b). The recent find of an anomalous, several-ton, in-situ, granitic boulder embedded in the Cleveland Shale Member of the Ohio Shale at the Cleveland-Bedford shale contact in northeastern Kentucky can only reflect a glacial-dropstone origin (Lierman and Mason, 2007) and clearly supports Late Devonian alpine glaciation in source areas to the east. This Late Devonian occurrence is the only evidence of Paleozoic glaciation in the Appalachian Basin, and according to Cecil et al. (2002, 2004), this time represents a period of cold, humid climate in the basin and coincides with a global, sea-level drawdown that contributed to unconformity development at the Devonian–Mississippian transition (Figure 4). The original eastward extent of the diamictites is uncertain because of erosion, but they must have been products of alpine and piedmont glaciation at the front of high Acadian mountains just formed at the New York promontory as a result third-tectophase transpression involving the Avalon and/or Carolina terranes (Figure 21B). Dennis (2005, 2007) has suggested that the diamictites sampled exotic Acadian highlands, probably formed during accretion of the Carolina terrane. Additional lithologic, stratigraphic, and chemostratigraphic evidence for such a glaciation is provided in Richardson and Ausich (2004). The distal ends of the coarser clastic parts of each cycle in the tectophase (Figure 24; i.e., upper OlentangyHanover-Angola, Chagrin, and Bedford-Berea) not only become thin and finer grained toward the western margin of the Appalachian Basin, but also become progressively blacker in color such that they are visually indistinguishable from the more proximal, black-shale parts of each cycle. Gamma-ray stratigraphy does, however,

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provide a way to distinguish the types of black-shale (Ettensohn et al., 1979; Ettensohn and Elam, 1985), and Ettensohn et al. (1988a) called them regressive and transgressive black shales, respectively (Figure 24). Hence, on the western margin of the Appalachian Basin in eastern Ohio, western West Virginia, eastern Kentucky, eastern Tennessee, western Virginia, northwestern Georgia, and northeastern Alabama where the coarser clastic parts of each cycle become black shales, the Catskill Delta complex grades westwardly into a nearly uniform, black-shale sequence, 1–570 m thick, known variously as the Ohio, New Albany, or Chattanooga shales (Ettensohn et al., 1988a) (Figure 24). This sequence is probably the most important oil and gas source rock in the Appalachian Basin and reflects infilling of the Acadian foreland basin, eustatic deepening, and the cratonward migration of Acadian subsidence beyond the foreland basin to form what Ettensohn (1998) has called a deep, cratonic, blackshale basin. This infilling of the foreland basin and the cratonward shift of the deeper water, black-shale facies to areas beyond the basin must reflect the greater extent and intensity of Acadian deformational loading during the third tectophase. The New England area bore the brunt of third-tectophase deformation, uplift, and erosion (Figure 21B), and hence, Late Devonian sedimentary deposits in the area are very rare. Nonetheless, small areas of volcanics, volcaniclastics, and other, commonly red, coarse, clastic sediments are known from the Wamsutta Formation of the Norfolk and Narragansett basins of Massachusetts and Rhode Island (Murray et al., 2001, 2004) (Figure 1), from the Perry Formation of the St. Andrews and Blacks Harbor basins of Maine and southern New Brunswick (Schluger, 1973) (Figure 1, along fault line cc), as well as from the Memramcook Formation of central New Brunswick, the Fleurant and Escuminac formations of northern New Brunswick and the Great Bay de l’Eau, Terrenceville and Pools Cove formations of Newfoundland (Poole et al., 1968); these formations have been interpreted to represent largely fluvial and alluvial-plain environments from high, intermontane, fault basins in the Acadian highlands. Evidence suggests that much of the sediment from these highlands was probably transported westwardly into parts of the Appalachian foreland basin that once overlay the Adirondack massif and areas to the north, suggesting that the Appalachian foreland basin in Devonian time continued uninterrupted northwardly into the Ontario Embayment, Quebec Basin, and perhaps beyond (Faill, 1997b). In fact, even though Devonian units are absent from these areas today, clasts of Devonian limestones and sandstones from the St. Helen’s Island Breccia in southern Quebec indicate that the Devonian foreland basin extended at least as far as the present Quebec Basin (Sanford, 1993b). The fourth tectophase of the Acadian orogeny (Ettensohn, 1985a) (Figure 4, cycle 11) is apparently synonymous with the so-called Neo-Acadian orogeny, first named by Robinson et al. (1998) based on Late Devonian–Early Mississippian magmatism, metamorphism, and deformation in central Massachusetts. Other evidence from the central Appalachians supports latest Devonian–earliest Mississippian dextral shear or accretion involving the Carolina terrane near the Virginia promontory at 35871 Ma (Boland and Dallmeyer, 1997; Hibbard et al., 2002; Trupe et al., 2003; Steltenpohl, 2005; Hatcher et al., 2005) (Figure 21B), latest Devonian–Late Mississippian (365–320 Ma) accretion of the Ordovician Dahlonega terrane and Hillabee back-arc volcanics at and near the Alabama promontory (McClellan et al., 2005a, b), and peak metamorphism in the central Appalachian Piedmont at about 360 Ma (Hatcher, 2005; Dennis, 2005, 2006a, b, 2007). Hatcher and Bream (2002), Hatcher et al. (2003, 2005), and Hatcher (2005) have also termed these coeval southern and central Appalachian events the Neo-Acadian orogeny, and it is likely that all of these Appalachian margin events were related and tied together via the Carolina terrane (Figure 21B). According to Dennis (2007), after initial Ordovician–Silurian docking (Figure 7), the Carolina terrane experienced renewed rifting, separation and transfer from Laurussia in Silurian–Early Devonian time during the Helderberg transition (Figure 21A); during this time, Carolina continued its northward, sinistral, transpressive dispersal until northern parts of the terrane docked with Laurussia in southern New England (Figure 21B). The Avalon terrane, however, perhaps with Meguma (van Staal, 2005), was caught in a pincers movement between Laurussia and Gondwana and had been translated obliquely southwestward since late Early Devonian time (Figure 21B). By Late Devonian time, the Avalon terrane had apparently collided with the north end of Carolina, explaining the intense deformation in Massachusetts noted by Robinson et al. (1998), and the oblique, dextral movement of Avalon terrane was transferred from this time onward to the Carolina terrane (Figure 21B). Today, the present boundary between the Carolina and adjacent Laurussian terranes is defined by a zone of dextral shearing and transpression (Boland and Dallmeyer, 1997; Hibbard et al., 2002; Trupe et al., 2003; Steltenpohl, 2005; Hatcher, 2005; Hatcher et al., 2005; Dennis, 2005; 2006a, b, 2007), especially in the area of the Virginia promontory. And in this area by the Late Devonian–Mississippian transition, dextral shearing and transpressional docking of the Carolina terrane against the Virginia promontory and points to the south were well under way (Figure 21B), explaining the intense deformational loading and uplift in adjacent highlands necessary to generate the rapid subsidence and eastward migration of the earliest Mississippian, Sunbury-Riddlesburg, black-shale foreland basin (Figures 21B and 24) and its later clastic infill (Figure 22). Dennis (2005; 2006a, b, 2007) has also suggested that vastly altered flysch in the Cat Square successor basin

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(Figure 21A), and perhaps other parts of the Inner Piedmont, were also displaced southwestwardly with the Carolina terrane. Hence, fourth-tectophase responses in the Appalachian foreland basin largely reflect Early Mississippian (Early Tournaisian; Kinderhookian) dextral transpression and shear at the Pennsylvania reentrant and Virginia promontory and the southward migration of orogeny toward the Alabama promontory during Early–Late Mississippian time (Figure 21B). The so-called Neo-Acadian orogeny is really nothing more than the fourth and final tectophase of north-to-south, zipper-like, transpressional convergence by peri-Gondwanan terranes suggested by Ettensohn (1985a, 1987) as the Acadian orogeny. Although this kind of transpressional scenario had been indicated earlier by the southern migration of Devonian–Mississippian black-shale units in the Appalachian Basin (Ettensohn, 1985a, 1987, 1994, 1998), more recent dating of plutonism, metamorphism, and arc collision from the area (e.g., Trupe et al., 2003; McClellan et al., 2005a, b; Hatcher et al., 2005) is helping to discern the specific tectonic mechanisms involved. The unusually long duration of this tectophase (Early Mississippian–Early Pennsylvanian: B45 Ma; Figure 4, cycle 11) may reflect the oblique nature of the convergence and the fact that the entire 650 km length of Caroline had to dock in this fashion (Figure 21B). The evidence for fourth tectophase cycle in the Appalachian Basin begins with a subtle Devonian– Mississippian unconformity behind each promontory (Ettensohn, 1994; Ettensohn and Pashin, 1997), overlain by the Sunbury black-shale and its equivalents (Figures 12, 14, and 24). Although the tripartite unconformity distribution behind each promontory clearly suggests tectonic influence, the unconformity in places also has the indented or bull’s-eye pattern (Ettensohn, 1994) characteristic of eustatic influence (Ettensohn and Pashin, 1997), and regional sea-level curves support such a drawdown in latest Devonian time (Dennison and Head, 1975; Dennison, 1989; Ettensohn and Pashin, 1997). The Sunbury is the most widespread of the Devonian–Mississippian black shales, but most of its distribution occurs west of and beyond the Appalachian Basin, marking an even more extensive development of the cratonic basin noted above (Figures 14 and 24). Hence in western and southwestern parts of the Appalachian Basin, Sunbury equivalents, if present, are included at the top of the New Albany and Chattanooga sequences (Ettensohn and Elam, 1985a; Ettensohn et al., 1988a). The Sunbury Shale is a deep, open-marine, fissile, black-shale in the western Appalachian Basin, but to the east in central parts of the basin, it grades into the brown-to-black, shallow- to marginal-marine Riddlesburg Shale (Figure 21B). The pattern in the flexural development of earlier, Devonian black-shale basins has been shown to be a progressive migration to the south and west (Figures 22 and 24), but the Sunbury-Riddlesburg black-shale basin is not only unusual in its extensive, westward, extrabasinal distribution, but also in the fact that its proximal basin boundary migrated eastward, rather than westward as in the other shale basins (Figures 12, 14, 21B, and 24), in some places by as much as 275 km. (Ettensohn, 1985a; Ettensohn et al., 1988a). This can only mean that the Neo-Acadian deformational loading responsible for Sunbury-Riddlesburg basin subsidence was somehow more intense, perhaps because of constriction to a narrow, more easterly belt. In fact, Hatcher (2005) has suggested that the buttressing effect of Blue Ridge crystalline rocks may have been important in intensifying and confining Neo-Acadian tectonism. In addition to its duration, the fourth tectophase is also unusual in the development of two distinct clastic wedges separated by a thick carbonate unit (Figure 14). The earliest clastic wedge, known as the Price-Pocono, Grainger, or Borden delta (Figures 14, 21B, 22, and 24), immediately followed basinal, Sunbury, black-shale deposition in south-central parts of the Appalachian Basin and represents loading-type relaxation following active tectonism. In west-central parts of the basin, however, dark-shale deposition equivalent to parts of this clastic wedge continued into Middle Mississippian (Tournaisian–Visean; Late Osagean) time despite declining sea levels at the time (Ross and Ross, 1987; Ettensohn et al., 2004b; Richardson and Ausich, 2004). These shales have been called ‘‘black Borden’’ (Ettensohn et al., 1988a; Figure 24) and apparently reflect continued basin subsidence accompanying Neo-Acadian deformational loading during peak Carolina convergence (e.g., McClellan et al., 2005a, b; Hatcher et al., 2005). Moreover, equivalent parts of the Price-Pocono largely reflect subaerial parts of the delta complex (e.g., Read, 1955; Kreisa and Bambach, 1973), whereas the Grainger and Borden represent more distal, subaqueous parts of the complex (e.g., Peterson and Kepferle, 1970; Rice et al., 1979), which include turbidites, prodelta, and delta-front deposits (Figures 22 and 24). The earliest and most distal parts of the delta in central and southern parts of the Appalachian Basin, and in some places the only part of the delta present, are thin, prodeltaic shale units like the New Providence and Maury shales, locally less than a meter thick. The delta complex has an unusually large distribution and an atypical paleoslope for this part of the succession; it extends 800 km along the length of the Appalachian Basin, overlapping large parts of the Catskill complex (Figure 22). Subaqueous lobes of the complex, however, extend cratonward nearly 600 km, crossing the Cincinnati arch and much of the Illinois Basin (Swan et al., 1965; Lineback, 1966) (Figure 22). The paleoslope was dominantly to the west, but in the Illinois Basin a southern paleoslope prevailed (Figure 22). The crossing of the Cincinnati arch and the southward turn into the Illinois Basin by the Borden delta complex is clearly unusual for a clastic wedge

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originating in the Appalachian area, but may reflect a nearly east-west belt of subsidence paralleling the front of concurrent Ouachita orogeny, which had begun on the southern Laurussian margin by latest Devonian–Early Mississippian time (Ross and Ross, 1985; Morton, 1985; Arbenz, 1989). Although the Borden may reach maximum thicknesses of 240 m in the western Appalachian Basin and 200 m in the Illinois Basin, proximal parts of the equivalent Price-Pocono complex (Figure 21B) attain maximum thicknesses of 500 m in the central Appalachian Basin (Meckel, 1970; Shaver, 1985; Patchen et al., 1985a). Relaxational, deltaic sedimentation continued in the Appalachian Basin until the Osage–Meramec (Early Visean) transition, when an abrupt shallowing occurred throughout the basin. Parts of this shallowing reflect a global, eustatic lowstand (Vail et al., 1977; Ross and Ross, 1987) (Figure 4), but the distribution of an accompanying regional unconformity also suggests uplift on a relaxational bulge that migrated toward the orogen (Ettensohn, 1994; Ettensohn and Pashin, 1997). Uplift of the bulge apparently cut off deltaic sedimentation, and in parts of the basin to the west and south, glaucony- and verdine-rich facies (see Odin, 1988; Thamban and Rao, 2000) in and around the widespread Floyds Knob Bed of the upper Borden and lower Ft. Payne formations (Stockdale, 1939; Kepferle, 1971; Sable and Dever, 1990; Ettensohn et al., 2004b) reflect extensive delta destruction and sediment starvation. In fact, from southeastern Kentucky and western Virginia to Georgia and Alabama, in the absence of major clastic influx, deeper water, cherty Ft. Payne carbonates infilled the basin (Gutschick and Sandberg, 1983; Ettensohn et al., 2002a), and at the Ft. Payne shelfbreak, bryozoan- and crinoid-rich, Waulsortian-like, carbonate mud mounds commonly developed (e.g., Ausich and Meyer, 1990); diagenetic alteration of these mounds created conditions conducive for subsequent petroleum accumulation (MacQuown and Perkins, 1982). At the same time in central parts of the basin on and near the bulge, deltaic sedimentation gave way to shallow, open-marine, peritidal, and evaporitic sedimentation in the Maccrady Formation (Warne, 1990). A subtle unconformity atop all the clastic-rich units in this same area (Figure 14) reflects the likely coincidence of bulge uplift and a lowstand maximum (Ettensohn, 1994; Ettensohn et al., 2002a, 2004a, b); this time of lowered sea level set the stage for a succeeding episode of major carbonate deposition throughout the basin. Delta destruction left behind a platform-to-ramp setting with little clastic influx, which, when combined with lowstand, shallow-water conditions and an evaporative, subtropical climate, produced ideal conditions for carbonate deposition. As a result, from near the Osage–Meramec transition to Mid-Chesterian (Early Visean– Early Serpukhovian) time in the Appalachian Basin, a period of about 18 Ma, widespread carbonate deposition supplanted the clastic infilling of the basin that would have normally characterized remaining parts of Acadian (Neo-Acadian) loading-type relaxation. These carbonates are present in units like the Tuscumbia, Monteagle, Bangor, Newman, Slade, Maxville, and Greenbrier limestones, which attain thicknesses ranging from 0 to 900 m (Figure 14). The carbonates reflect environments varying from deep-ramp to exposed paleosols during two thirdorder eustatic cycles, the effects of which have been locally altered by synsedimentary tectonism (Ettensohn 1980, 1981; Ettensohn et al., 1988b, 2004a, b) and fourth-order, shallowing-upward cycles that are thought to reflect glacio-eustatic fluctuations during the early stages of Carboniferous–Permian, Gondwanan glaciation (Al-Tawil et al., 2003; Al-Tawil and Read, 2003; Ettensohn et al., 2004b). By Late Chesterian (Early Serpukhovian) time, carbonate deposition was abruptly interrupted by a brief period of deeper water, shale deposition, reflected in units like the informal ‘‘Pencil Cave’’ shale, upper Newman or Bluefield formations, and the Lillydale or Maddox Branch shales, which probably represent infilling of the peripheral sag accompanying unloading-type relaxation (Figure 7). With unloading-type relaxation came a cratonward-prograding wedge of largely Upper Mississippian (Serpukhovian; Late Chesterian) marginal-marine and terrestrial clastic sediments with abundant redbeds in the Pennington and Mauch Chunk groups (Figure 22). These units extend approximately 1,000 km along the length of the central and southern Appalachian Basin (Figure 22), attaining a maximum thickness of 2,700 m. (Patchen et al., 1985b; Milici and de Witt, 1988); their great extent and thickness may in part be related to fact that by Late Chesterian time, the Appalachian Basin had moved into the humid, equatorial zone where clastic deposition is favored (Scotese, 2003; Cecil et al., 2004) (Figure 4). Upper parts of these two units may locally be gradational with or contain sediments of earliest Pennsylvanian (Bashkirian; Morrowan) age in the Pocahontas Formation and equivalents (Englund et al., 1979; Patchen et al., 1985b). Sequence deposition was ended by widespread erosional truncation on the Early Pennsylvanian, sub-Absaroka unconformity (Ettensohn and Chesnut, 1989; Ettensohn, 1994). Although many workers have suggested that this thick Upper Mississippian clastic wedge reflects inception of the Alleghanian orogeny (e.g., Davis and Ehrlich, 1974; Perry, 1978; Milici and de Witt, 1988; Chesnut, 1991; Hatcher, 2005), flexural indicators for inception of orogeny are absent: namely, presence of a bounding basal unconformity, a basal transgressive sequence, and evidence of two opposed paleoslopes. In fact, absence of the above criteria, as well as the predominance of marginal-marine facies with redbeds and evidence for a dominant cratonward-dipping paleoslope (e.g., Colton, 1970; Meckel, 1970), suggest that unloading-type relaxation continued into earliest Pennsylvanian time. Unloading-type rebound, however, may have been accompanied by

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uplift related to continued assembly of Neo-Acadian (McClellan et al., 2005a, b; Hatcher, 2005; Hatcher et al., 2005), or possibly Early Alleghanian (Goldberg and Dallmeyer, 1997), thrusts in the subsurface as early as Late Mississippian (Late Chesterian; Serpukhovian) time, helping to explain the early, and extensive, development of the Mauch Chunk-Pennington clastic wedge, even as carbonates predominated nearly everywhere else in the basin. The presence of smaller scale sequences throughout the Pennington and its equivalents indicates the continuation of Gondwanan glacio-eustatic cycles into Late Chesterian (Serpukhovian) time (Miller and Eriksson, 1999) (Figure 4). 9.2.2. Influence of the Ouachita orogeny (latest Devonian–Early Pennsylvanian, B360–B312 Ma) The Ouachita orogeny probably reflects convergence between one or more of the microcontinents or oceanic plateaus that preceded Gondwana, and the southern margin of Laurussia (Lowe, 1985; Arbenz, 1989; see Miall, Chapter 8). Although evidence suggests rising tectonic highlands to the south and southeast of Laurussia by Late Devonian and earliest Mississippian time (Ross and Ross, 1985; Morton, 1985), docking with southern parts of Laurussia probably did not occur until Late Mississippian (Early Chesterian; Late Visean) time (Thomas, 1989). When docking did occur, it apparently began at the Alabama promontory and proceeded westward in time (Thomas, 1989). Accompanying synorogenic sedimentation was largely restricted to the resulting Black Warrior peripheral basin (Figure 1), but there is evidence for flexural interaction with southernmost parts of the Appalachian Basin (Ettensohn and Pashin, 1993) and for Ouachita-related unconformities that overlapped into the Appalachian Basin (Ettensohn, 1993, 1994; Ettensohn and Pashin, 1997). The oldest unconformity occurs at the Kinderhook-Osage (Mid-Tournaisian) transition and parallels the Ouachita orogen. Ettensohn (1994) suggested that it probably reflects bulge moveout during an Early Ouachita tectophase marking inception of Ouachita convergence in the area; related uplift may have influenced the transition from deeper water Ft. Payne cherts to shallow-water carbonates in southernmost parts of the Appalachian Basin (Ettensohn, 1993). The second major unconformity is an Early Chesterian surface at the Genevievian–Gasperian (Late Visean) transition that occurs across southern and southwestern parts of the Appalachian Basin. The facts that unconformity distribution again parallels the Ouachita orogen (Ettensohn, 1993, 1994) and that unconformity age and subsequent infilling of the Black Warrior Basin coincide with the Late Mississippian inception of docking at the promontory (Thomas, 1989; Thomas and Whiting, 1995) suggest that the unconformity was related to Ouachita bulge moveout (Ettensohn and Pashin, 1993, 1997), although eustatic lowstand at the time (Ross and Ross, 1987, 1988; Dennison, 1989) (Figure 4) may have also been contributory. Finally, Figure 25 suggests how parts of the previously discussed Appalachian (Acadian) relaxational sequence (Monteagle–Pennington formations) were modified by Ouachita flexural influence in southwestern parts of the Appalachian Basin. Most notably, the Bangor carbonate bank thickens substantially on the probable Ouachita bulge in the area of the Black Warrior–Appalachian transition and tongues out into back-bulge, relaxational (loading-type), foreland-basin, clastic infill; the northeasterly sourced Pennington Formation apparently pinches out in front of the bulge due to the growing Bangor bank; and the Evans and Hartselle sandstones reflect prograding siliciclastic rims around the southern margin of the Monteagle carbonate bank (Cleaves and Broussard, 1980), which represents an earlier bulge position (Figure 25). Further information about flexural interactions between the Black Warrior and Appalachian basins is presented in Ettensohn and Pashin (1993). 9.2.3. Alleghanian orogeny and foreland-basin sedimentation (Early Pennsylvanian–Permian?; 315–265?) During the Alleghanian orogeny, the Appalachian Basin and the adjacent Alleghanian highlands were largely situated in the tropical, equatorial belt (Scotese, 2003), characterized by long-term, humid to perhumid climate (Cecil et al., 2003, 2004). At nearly the same time, global icehouse conditions had begun at high latitudes on Gondwana and culminated during Permian time (Fischer, 1984; Frakes et al., 1992). Sea levels overall were low (Fischer, 1984; Ross and Ross, 1987; Dennison, 1989), but a pronounced fourth- and fifth-order cyclicity related to glacio-eustasy, climate, and tectonics (Figures 4) seems to have controlled sedimentation throughout most of the Appalachian Basin (Cecil, 1990; Chesnut, 1994; Cecil et al., 2003, 2004), and was the major factor in development of the classic cyclothems (see Miall, Chapter 8). The Alleghanian orogeny was the last Appalachian compressional phase and represents final closure of the Rheic Ocean during the zipper-like amalgamation of Gondwana (Africa) and Laurussia. Much of the orogeny probably reflects oblique convergence, and in some parts of the orogen, it was a largely transpressional event with more strike-slip than compressional deformation (e.g., Gates et al., 1986; Hatcher, 1999, 2002, 2005; see Gibling et al., Chapter 6). Although there is evidence of dextral shear and crustal-escape mechanisms, the central and southern Appalachians reflect a dominantly convergent tectonic zone, mainly characterized by a broad fold-thrust

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Figure 25 Schematic southwest-northeast section from the Black Warrior Basin into the southern end of the Appalachian Basin (see Figure 1), showing the transition between basins and the overlap of Ouachita unconformities and units into the Appalachian Basin. The Lewis-through-Parkwood section is an Ouachita £exural sequence similar to those in the Appalachian Basin. The positions of the main bodies of the Monteagle and Bangor limestones represent the successive positions of Ouachita bulges (from Ettensohn and Pashin, 1993).

belt; in particular, the Blue Ridge–Piedmont megathrust transported new crust formed in previous Paleozoic orogenies at least 350 km westward across the Laurentian platform margin and parts of the Appalachian foreland basin (Figure 2) (e.g., Hatcher, 2005). In contrast, the northern Appalachians reflect more complex wrench and escape tectonics, but structural domains in both parts of the Appalachians seem to have been connected via dextral shear zones that corresponded to older Precambrian faults formed during earlier Iapetan rifting (Arthaud and Matte, 1977; Getty and Gromet, 1988; Manspeizer, 1994). The inception of convergence is poorly constrained, but if the major sub-Absaroka or Monday Creek unconformity (Sloss, 1963) is interpreted to represent the inception of orogeny, then based on evidence from the Appalachian Basin, Alleghanian orogeny began in late Early Pennsylvanian (Bashkirian; Morrowan) time (Ettensohn and Chesnut, 1989; Ettensohn, 1994). The clockwise convergence of Gondwana toward Laurussia (e.g., Ziegler, 1989; Scotese, 1998; Hatcher, 2002, 2005), structural evidence (Rodgers, 1967; Dean et al., 1988), age and distribution of clastic wedges in the foreland basin (Chesnut, 1989; Patchen et al., 1985a, b; Hatcher, 2005), and flexural modeling (Beaumont et al., 1987, 1988) all suggest that the orogeny proceeded from south to north. However, this timing and progression are difficult to reconcile with the possibility that Gondwanan convergence may have begun in the Maritime Canadian region as early as Late Devonian time (Ziegler, 1989) (Figure 21B), and certainly by Late Mississippian time in the northern Appalachian area (Hatcher, 2002, 2005). This seeming contradiction, however, may merely reflect the fact that the Alleghanian orogeny was not mainly a subduction-type orogeny, but more of a collisional orogeny with an extensive, transpressional zone related to escape tectonism during the convergence and rotation of Gondwana against the southeastern margin of Laurussia (e.g., Gates et al., 1986, 1988; Hatcher, 2002, 2005). Actual docking between the Reguibat promontory of Africa on Gondwana (Figure 21B) and the peri-Gondwanan collage of microcontinents accreted to Laurussia took place at about 315 Ma near the New York promontory (Faill, 1997b; Hatcher, 2002, 2005; Engelder and Whitaker, 2006) and only continued the already prominent, dextral, strike-slip tectonics that had begun with Avalonian and Carolina terranes during the Acadian orogeny. Dextral shear and escape tectonics were prominent along all parts

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of the orogen (Snoke et al., 1980; Mosher, 1983; Gates and Glover, 1989; Hatcher, 2001), but in early phases of the orogeny may not have formed highlands significant enough to drain into the Appalachian foreland basin, or sedimentation may have been restricted to transtensional basins along shear zones like the Narragansett Basin in Rhode Island and Massachusetts (e.g., Mosher, 1983; Murray et al., 2004) (Figure 1). By latest Early Pennsylvanian time, however, Gondwana began to pivot clockwise around the New York promontory, which initiated head-on collision with southeastern Laurussia and emplacement of the Blue Ridge–Piedmont megathrust in the central and southern Appalachians (Hatcher, 2002, 2005). At this time, many earlier strike-slip fabrics were overprinted by brittle, dip-slip motion (Hatcher, 2001), and evidence for the orogeny began to accumulate in the Appalachian foreland basin. A major unconformity formed and the earliest sediments entered the basin from the south. These sediments also record arrival of the first Paleozoic detrital zircons in the Appalachian foreland basin (Hatcher, 2005), indicating erosion of crust formed during earlier Paleozoic orogenies. The resulting Allegheny clastic wedge extends more than 1,300 km along the length of the Black Warrior– Appalachian foreland basin (Figure 22). The sedimentary record of this orogeny, moreover, is very unlike that produced during previous tectophase cycles and is composed almost entirely of terrestrial and marginal-marine, molasse-like sediments; deeper water, dark, marine shales and associated flysch-like sediments are largely absent. The thickest accumulations of these siliciclastic sediments, up to 2,900 m thick, occur in the foreland basin just cratonward of respective promontories (Meckel, 1967; Colton, 1970; Cook and Bally, 1975; Patchen et al., 1985a, b). Unlike the distribution of clastic wedges in previous orogenies, a blanket of siliciclastic sediment advanced westward more than 1,000 km across large parts of the east-central Laurussian craton (Figure 22), reflecting an overfilled foreland basin (see Flemings and Jordan, 1989; Jordan, 1995). Across much of the adjacent craton and throughout most of the foreland basin, this clastic blanket is partitioned into a hierarchy of cycles by thin marine zones or coals (Figure 26) that reflect a combination of climatic, glacio-eustatic, and tectonic control (e.g., Chesnut, 1994; Heckel, 1994, 2002; Cecil et al., 2003, 2004; see Miall, Chapter 8). Although many deltaic models have been proposed to explain the origin and distribution of these sediments (e.g., Ferm and Cavaroc, 1969; Wanless et al., 1970; Ferm, 1970, 1974; Donaldson and Shumaker, 1981), deltas were probably only minor parts of the depositional setting; various paralic, estuarine, fluvial, and alluvial-plain environments were apparently more common. Throughout most of the Appalachian Basin, Mississippian and Pennsylvanian rocks are separated by the so-called Mississippian–Pennsylvanian, sub-Absaroka (Sloss, 1963), or Monday Creek (Dennison, 1989) unconformity. Despite its Mississippian–Pennsylvanian appellation, it is largely Early Pennsylvanian in age (Englund et al., 1979; Chesnut, 1989, 1992; Pashin et al., 1991). With the post-Sauk unconformity (Figure 16), it is one of the most extensive unconformities in the Appalachian Basin, and begins as an intra-Pennsylvanian unconformity in the east and progressively truncates older Mississippian rocks to the west (Meckel, 1970; Ettensohn, 1994). Locally, in areas just behind the three promontories, the unconformity is probably absent, a situation that must reflect major subsidence that far outstripped uplift because of increased deformational loading at the promontories (Ettensohn, 1994). Although in places the unconformity may reflect components of a latest Mississippian eustatic lowstand (Saunders and Ramsbottom, 1986; Ross and Ross, 1987, 1988) (Figure 1), in large part, the surface more likely reflects bulge uplift and migration or continental braking accompanying inception of the Alleghanian orogeny (Quinlan and Beaumont, 1984; Beaumont et al., 1987, 1988; Ettensohn and Chesnut, 1989; see Burgess, Chapter 2). It seems that early shear and escape phases of the orogeny were apparently not sufficient to generate significant uplift in the foreland; only with the head-on collision in central and southern parts of the orogen was loading significant enough to generate foreland bulge moveout and uplift. The nearly continent-wide extent of the unconformity (e.g., Ham and Wilson, 1967), however, suggests that a major component of continental braking related to Gondwana–Laurussia collision or concurrent Ouachita orogeny in the south may have also been involved. Concurrent with the initiation of head-on-collision in the south, possible arc development (Sinha and Zietz, 1982; Dallmeyer, 1986) and unconformity development (Englund et al., 1979; Chesnut, 1989, 1992; Pashin et al., 1991), the effects of the Alleghanian orogeny were first apparent in the foreland basin by late Early Pennsylvanian (Late Namurian or Bashkirian; Morrowan) time. Despite some understanding about the timing and progression of orogeny, at least as indicated from the foreland basin, interpretation of the sedimentary record regarding the manifestation of tectophase cycles is still uncertain. Based on structural criteria, Geiser and Engelder (1983) suggested that there were two Alleghanian tectophases, the earliest of which they called the Lackawana tectophase. In contrast, Donaldson and Eble (1991) suggested the presence of three tectophases, but their earliest tectophase includes lowest parts of the Lower Pennsylvanian section that are gradational with underlying Mississippian rocks, and hence, may reflect final Acadian or Neo-Acadian relaxation (Ettensohn and Chesnut, 1989; Ettensohn, 1994, 2004) with some possible overlap by Early Alleghanian or Ouachita assembly; their final two tectophases, however, are divided at nearly the same point suggested later in this chapter. Miller and Kent (1988) also suggested two major thrusting events,

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Figure 26 Schematic, west-east, sectional diagram showing the eight, fourth-order, Lower Pennsylvanian cycles of the Breathitt Group, which largely comprise the Lackawana tectophase. Each cycle begins with a named marine zone that represents marine £ooding of the foreland basin. The basin then in¢lled with ¢ve to seven, ¢fth-order, coal-clastic cycles of likely glacio-eustatic origin represented on the right by successive coal beds (undulating lines with rooting). In the lower four cycles, the coal-clastic cycles were truncated by quartzarenite, sandbelt complexes representing braidplains that migrated westward in time with each new cycle. The arrowheads in the four lower cycles point to mid-formational, marginal-marine shales that may re£ect further subdivision of the lower four cycles. The trailing, loose stipple at the eastern termination of each sandstone body represents the possibility of some eastern, transverse streams feeding the braidplain channels.While the lower four cycles probably represent ‘‘loading-type’’ relaxation, the upper four cycles, lacking major quartzarenite channel sands, apparently prograded beyond the limits of the former foreland basin and probably represent ‘‘unloading-type’’ relaxation.The overlying, Upper Pennsylvanian Conemaugh and Monongahela formations re£ect a new tectophase and may be separated from the Princess Formation by a subtle unconformity. The diagonally lined area at the base of the diagram is the Early Pennsylvanian, sub-Absaroka or ‘‘Mississippian--Pennsylvanian’’ unconformity. Adapted from Chesnut (1992).

an earlier event in the southern Appalachians and a later one to the north, events that may also reflect two tectophases. The Lackawana tectophase may reflect early development of west-directed, oblique subduction below the newly accreted southeastern margin of Laurussia and development of a magmatic arc forward of the Alabama and Virginia promontories during Early and Middle Pennsylvanian time (Bashkirian–Moscovian; Morrowan– Desmoinesian) (Sinha and Zietz, 1982; Dallmeyer, 1986), although Samson et al. (1995) have indicated that subduction was unlikely. Nonetheless, concomitant loading, largely through thrust-de´collement tectonism to the south (Faill, 1998), developed a foreland basin throughout southern and central parts of the larger Appalachian Basin. The extent of individual units in that basin and the very shallow-marine to terrestrial nature of included sediments suggest that the basin was broad and shallow, perhaps reflecting increased crustal rigidity and the fact that deformation had advanced more expansively cratonward than during previous orogenies. Sediments deposited during the tectophase are included in the Breathitt Group (Chesnut, 1992, 1994, 1996; Aitken and Flint, 1995) (Figure 26) and exhibit a pronounced, fourth-order, transgressive-regressive cyclicity of thin marine horizons with overlying coal-bearing sequences of largely estuarine, tidal, or alluvial-plain origin, which each containing groupings of five to seven, fifth-order, coal-clastic cycles (Greb et al., 2004) (Figure 26). The fourthorder cycles are defined by marine-flooding surfaces at their bases and, based on interpolation from radiometric dating, reflect B2.5 Ma of time; they were probably tectonic in origin and include the fifth-order cycles with B0.4 Ma periodicities of likely glacio-eustatic origin (e.g., Chesnut, 1992, 1994, 1996) (Figure 26). In the four earliest of the fourth-order cycles, the cratonward margin of each coal-bearing sequence is truncated by a channel

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complex of quartz-pebble-bearing quartzarenites (Warren Point, Sewanee, Bee Rock, and Corbin sandstones; Figure 26), interpreted to represent braided-stream sandbelts that ran longitudinally along the distal margin of the foreland basin on a south- to southwest-dipping paleoslope (Rice and Schwietering, 1988; Chesnut, 1994; Greb and Chesnut, 1996; Greb et al., 2002, 2004). Each of the sandstones appears to be further divided by a midformational shale, which also seems to reflect a marine-flooding surface that probably divides each of the cycles containing distal quartzarenites into two fifth-order cycles (Greb et al., 2004) (Figure 26, arrowheads). The thicker, more basinward, coal-bearing parts of each fourth-order cycle are characterized by litharenites and sublitharenites in channels that trend westward or northwestward, transverse to the foreland basin and to the longitudinally oriented quartzarenite sandbelts. The quartzarenite sandbelts appear to truncate most of the transverse channels and adjacent parts of the coal-clastic cycles containing them. The exception probably occurred at the top of each cycle below the next marine-flooding event, where the transverse channels must have been gradational into the longitudinal sandbelt and provided at least some of the quartz sands and pebbles for the longitudinal belts; this exception implies that there must have been different sources for transverse streams in the earlier coal-bearing parts of the cycle than for the distal, longitudinal streams and their feeder streams in later parts of each cycle (Greb et al., 2004). The western or cratonward boundary of each longitudinal sandbelt was apparently the peripheral bulge (Chesnut, 1994), and the fact that the bulge, sandbelts, and related coal measures shifted progressively westward in time and space (Figure 26) suggests a likely tectonic origin in response to a shifting deformational load (Greb et al., 2004). As has been suggested for similarly shifting Devonian– Mississippian black-shale subcycles (Figures 12, 23, and 24), each of these fourth-order cycles may reflect the cratonward movement of a major thrust system (Tankard, 1986; Greb et al., 2004). Overall, the lower four, Lower Pennsylvanian (Bashkirian; Morrowan) cycles are unique in the presence of shifting quartzarenite belts and in the less marine, more estuarine nature of their truncated, basinward, coal-bearing parts (Figure 26). These belts represent an Amazon-scale drainage system that headed in the Canadian shield with a paleodrainage that ranged from 1.3 106 to 2.9 106 km2 (Archer and Greb, 1995; Gray and Zeitler, 1996). At the southern end of the Appalachian Basin in present-day Alabama and Georgia, the fluvial nature of these sands changes to beach and barrier-island facies (Ferm and Ehrlich, 1967; Hobday, 1974), reflecting transportation and reworking of the sands in a marine embayment (Figure 27). The overlying, Middle Pennsylvanian (Moscovian; Atokan–Desmoinesian), fourth-order cycles are similarly bound by regionally extensive, marine zones, but they lack the distal, truncating, quartzarenitic, channel sands, suggesting the absence of a major, quartz-sand, sediment source. In addition, the marine zones are much more

Figure 27 Schematic paleogeographic diagram of the collision zone between Gondwana and the southeastern margin of Laurussia during Early Pennsylvanian time.The compact stipple represents one of the braidplain sandbelts running longitudinally along the northwest margin of the foreland basin; this sandbelt has truncated some of the earlier, transverse drainage between it and the thrust belt to the southeast. As suggested in the ¢gure, most of the major drainage (loose stipple) from the thrust belt probably entered the braidplain from the north and northeast. During early parts of each fourth-order cycle (see Figure 26), the foreland basin (braidplain and transverse drainage) would have been £ooded with marine waters from the south forming an elongate estuary/tidal complex (from Ettensohn, 2005).

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extensive, and the intervening coal-clastic cycles are more tidal in origin up to the level of the peak, marineflooding event (Magoffin Member), after which the cycles become more fluvially dominated; each of these cycles, moreover, apparently prograded beyond the limits of the foreland basin (Figure 26), indicating an overfilled (see Flemings and Jordan, 1989; Jordan, 1995) foreland basin (Chesnut, 1994; Greb et al., 2004). The overlying Upper Pennsylvanian (Kasimovian; Missourian) units of the Conemaugh Group may unconformably overlie the Middle Pennsylvanian units just discussed and apparently represent a different depositional framework with different sources in a different basin, and hence, they probably reflect a new tectophase. Clearly, the sedimentary sequence and lithologies present in the Lower–Middle Pennsylvanian Breathitt Group appear to be vastly different than those that characterize earlier tectophases (e.g., Figures 9 and 12). These differences no doubt reflect the collisional nature of the Alleghanian orogeny and a resulting broad, shallow foreland basin, cyclic glacio-eustatic influences, and presence in the humid equatorial belt (e.g., Cecil and DuLong, 2003; Cecil et al., 2003; Cecil et al., 2004) (Figures 4 and 27). Nevertheless, elements of the Breathitt sequence still broadly parallel typical tectophase development, and suggest that the entire sequence does represent a distinct tectophase (Figure 4, cycle 12). The tectophase sequence begins with the major sub-Absaroka unconformity (Figures 4, 14, and 26), reflecting at least in part major bulge uplift and moveout. Moreover, if each of the four, fourth-order (or eight fifth-order) Early Pennsylvanian cycles with distal quartzarenites is interpreted to represent cratonward movement of a major thrust system, then each of the cratonward-shifting, quartzarenitebearing, cycle sequences (Figure 26) is similar to Devonian–Mississippian black-shale subcycles (Figures 12 and 24). The marine-flooding event at the base of each cycle, which is represented by dark shales and/or dark limestones, apparently reflects basin subsidence accompanying thrust moveout and loading. The overlying series of coal-clastic cycles (Figure 26; Pikeville–Princess formations) probably represents loading-type relaxation with the westerly or southwesterly directed marginal-marine and terrestrial clastic sediments taking the place of flyschlike clastic sediments in a shallow foreland basin. This part of the section also contains bentonites (Chesnut, 1985; Rice et al., 1994). The smaller fifth- or sixth-order, coal-clastic cycles that nearly fill each basin at this stage are merely responses to climatic and eustatic controls caused by alternating glacial (coal-rich) and interglacial (clasticrich) episodes (Cecil and DuLong, 2003; Cecil et al., 2003, 2004). The muds, litharenites and sublitharenites that predominate in this phase of each cycle largely represent the sedimentary unroofing of each thrust complex and probably filled, and even overfilled, the shallow foreland basin, especially toward its eastern margin, which was proximal to source areas; and indeed, all cycles do thicken in this direction (Greb et al., 2004). This pattern of filling or overfilling toward eastern source areas may have also been reinforced by slow thrust velocity (Flemings and Jordan, 1989) or by climatic asymmetry and a resulting high erosive efficiency, inasmuch as the Pennsylvanian paleoequator and west-directed trade winds that encroached upon it from the southeast and transported weather systems toward the basin, were aligned on the foreland-basin side of the Alleghanian highlands (e.g., Scotese, 1998; Cecil et al., 2003) (Figure 27). With low westward gradients into the nearly filled foreland basin, the thrust complexes, now unroofed to the level of Early Paleozoic sandstones or basement igneous and metamorphic rocks, would have experienced higher gradients to the east, northeast, and north and generated high-gradient streams carrying quartz-rich sediments in those directions (Figure 27). Some of the transverse, westward drainage may have also transported quartz-rich sediments westward, and there is evidence for this in later parts of each cycle, but most paleocurrent evidence from the quartzarenite sandbelt complexes supports transport from the north and northeast (Greb et al., 2004). Because continental collision effectively prevented the symmetrical development of major sedimentary sinks to the east and southeast, much sediment generated on the eastern flanks of the mountains must have emerged from the mountains toward lower areas in the north that had not yet experienced major Alleghanian convergence. These streams emerged from the mountains in the north and northeast with their quartz-rich bedloads to form the large, low-gradient, braided, trunk streams that then flowed southward along the lowest, western margin of the foreland basin toward the deepest, most rapidly subsiding parts of the foreland basin in what is now Alabama (Figure 27). The peripheral bulge no doubt acted as a westward limit for the braided-stream complexes (Chesnut, 1994), but as successive thrust-sheet loads migrated westward, so did the bulge and successive braided-stream belts (Figure 26). Cratonward migration of the braided-stream sandbelts occurred at least four, and possibly eight, times before the thrust complexes ceased moving and were weathered and eroded to near elevational equilibrium with the filled foreland basin, as indicated by the absence of major quartzarenite units in subsequent parts of the Breathitt Group. These segregated quartzarenite units would seem to be unusual manifestations of loading-type relaxation, even under a collisional regime, but in the Appalachian Basin they probably represent the unique coincidence of climatic and tectonic asymmetry (e.g., Johnson and Beaumont, 1995), and their absence in overlying units probably signals the beginning of unloading-type relaxation near the end of the tectophase. By Middle Pennsylvanian time, distal, sandbelt development had ceased, the foreland basin had filled, and four, successive, fourth-order cycles defined by major marine zones (Pikeville, Hyden, Four Corners, and Princess formations; Figure 26) apparently overflowed cratonward beyond the basin (Chesnut, 1994;

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Greb et al., 2004). The lower three cycles are defined by marine zones (Betsie, Kendrick, and Magoffin members; Figure 26) that probably represent the most widespread, Pennsylvanian, marine-flooding events in the Appalachian Basin (Greb et al., 2002). In the two, intervening, coal-clastic sequences, coals are widespread, and the immature clastic sediments therein are tidally dominated and largely reflect a west-dipping paleoslope (Chesnut, 1994; Greb et al., 2004). The two overlying cycles, defined respectively by the marine, Magoffin, and Stoney Fork members at their bases, are more sandy and fluvial-dominated, and in the final cycle, the direction of thickening begins to change from southeast into the Central Appalachian Basin, toward the east and north into the Northern Appalachian or Dunkard Basin (Figure 28).

Figure 28 Map of east-central United States showing general locations of the Central Appalachian and Dunkard (Northern Appalachian) basins, which in large part represent foreland basins for the Lackawana and second Alleghanian tectophases, respectively. Large dark arrows represent major paleoslope directions in each basin, whereas the smaller arrowheads represent secondary directions. The northward migration of foreland basins in time apparently tracks the progress of the Alleghanian orogeny after collision in the south.

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Although continued presence of fourth-order cycles in the Middle Pennsylvanian section seems to reflect some continuation of cyclic loading, the absence of major quartzarenite-sandbelt sedimentation probably reflects the lowering of major source areas and the establishment some degree of elevational equilibrium between these areas and the largely filled foreland basin, which marks the onset of unloading-type relaxation. Transition to this stage of loading is also supported by the predominance of a west-dipping paleoslope and apparent overflow of the filled foreland basin. The widespread nature of the three earliest Middle Pennsylvanian marine zones, moreover, may indicate influence by the anti-peripheral bulge or peripheral sag described earlier, during which thin marine horizons become very widespread (Ettensohn, 1994). Hence, Lower and Middle Pennsylvanian units of the southern and central Appalachian Basin seem to be best described in terms of a single, foreland-basin, tectophase cycle. The general pattern of the sequence cycle still appears to be present, but it is overprinted by the effects of a shallow foreland basin, which insured the predominance of marginal-marine to terrestrial sedimentation in every part of the cycle; by the effects of tropical climatic asymmetry, which forced the predominance of clastic or organic sediments in every part of the cycle; and by the effects of glacio-eustatic and tectonic cycles, which imparted other levels of cyclicity across all parts of the larger tectophase cycle. Also interesting is the absence of Alleghanian-age detrital zircons in these sediments, which indicates that Alleghanian igneous rocks had not yet been exhumed or integrated into the drainage at this time (Thomas et al., 2004). By Late Pennsylvanian, and possibly Early Permian (Kasimovian–Artinskian; Missourian–Leonardian), time, foreland-basin sedimentation shifted northward into the northern Appalachian or Dunkard Basin (Figure 28), apparently tracking the similar progression of Alleghanian orogeny. The sequence of resulting sediments, comprising the Conemaugh, Monongahela, and Dunkard groups, may unconformably overlie the Middle Pennsylvanian Breathitt Group and its equivalents and probably represents a second and final Alleghanian tectophase (Figure 4, cycle 13); loading apparently reflected a change in structural style to fold-and-thrustde´collement tectonism in adjacent parts of the Alleghanian orogen and to non-de´collement, wrench tectonism in the northern Appalachians (Faill, 1998). In contrast to deposition during the Early and Middle Pennsylvanian Lackawana tectophase, in which marine flooding entered the southern and central Appalachian foreland basin from the south (Greb and Chesnut, 1996), during early parts of the final tectophase, subsidence in the north opened the Dunkard foreland basin to brief episodes of marine flooding from the Illinois Basin, perhaps through structurally related sags in the Cincinnati arch like the Sebree trough and Cumberland saddle (see Chesnut, 1994; Ettensohn et al., 2002b). By this time, Ouachita and Early Alleghanian orogenies had closed seaways to the south (see Miall, Chapter 8), and deformational loading to the north during the new tectophase apparently lowered the Cincinnati Arch and focused foreland-basin subsidence in the area of the Dunkard Basin (Figure 28). The tectophase began with exposure along a possible unconformity and subsequent marine inundation, represented by the Brush Creek Limestone (Figure 26). The Brush Creek, together with five, overlying, cyclic, marine horizons, may represent an early period of overall transgression from the west in response to cyclic thrust loading in the orogen. By Late Conemaugh time, deltaic and alluvial-plain sediments, prograding from sources around the basin margin during loading-type relaxation (Figure 28) effectively closed the basin to further marine inundation and generated a period of starved-basin infilling by chemical (lacustrine limestone) and organic (coal) deposits in large fresh-water lakes and swamps, represented by the Monongahela Group (Kovach, 1979; Donaldson and Shumaker, 1981; Donaldson and Eble, 1991). The presence of distinct basin geometry with multiple paleoslopes suggests continuation of loading-type relaxation. By latest Pennsylvanian and possibly Early Permian time, the basin had apparently filled with alluvial-plain sediments of the Dunkard Group that prograded into the basin from the south and east (Kovach, 1979; Donaldson and Shumaker, 1981; Milici and de Witt, 1988). This progradation probably marked inception of unloading-type relaxation and was the final phase of Appalachian sedimentation. The foreland-basin tectophase cycle manifest by this tectophase is even less distinct than the previous one, but still shows early transgression followed by two different episodes of relaxation-related progradation. Although cyclicity is present, larger fourthorder cycles are absent or indistinct, perhaps reflecting the fact that the final Alleghanian tectophase was largely the product of wrench or strike-slip, escape tectonism that tightened the fit between Gondwana, Laurussia, and intervening peri-Gondwanan terranes to the north (e.g., Gates et al., 1986, 1988; Manspeizer, 1994; see Chapter 6). The result of this wrench movement in New England is a series of small, largely non-marine, alluvial basins in Rhode Island, Massachusetts, and Maine that developed as pull-apart basins along some of the strike-slip faults (Bradley, 1982; Mosher, 1983; Hatcher et al., 1989; Murray et al., 2004; Gibling et al., Chapter 6) (Figure 1, fault cc). In the Narragansett Basin of Rhode Island and Massachusetts, the infill of sandstone, conglomerate, and coal may approach 4,000 m in thickness (Skehan, 1985), but in this basin, at least, sediment had been accumulating since Late Devonian time (Murray et al., 2004).

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10. Alleghanian Mountains, Post-Orogenic Collapse and Extension By the end of Permian time, the Alleghanian orogen was probably a high-standing, Himalayan-type mountain range with an added orogenic thickening of 3–10 km based on structural and thermal modeling (Levine, 1986; Friedman, 1994; Faill, 1998). The mountains may have included a broad altiplano across much of the southern and central Appalachians (Rodgers, 1987) that would have been similar to the modern central Andes, a 300 km-wide range with a mean elevation of about 4 km (Beaumont et al., 1987; Slingerland and Furlong, 1989; Faill, 1998). Preserved sediment in the foreland, moreover, seems to reflect sediments of older, recycled, orogenic provenance, largely from earlier Alleghanian tectonism to the south (Faill, 1998; Thomas et al., 2004). In the north-central Appalachians, however, the presence of units representing relief-sensitive estuarine and lacustrine environments as far east as northeastern Pennsylvania suggests that latest Alleghanian deformation was post-clastic wedge and post-Early Permian in age. Even though parts of the Appalachian Basin itself must have experienced major faulting and folding after this time, deformation may not have been readily visible, as the basin almost certainly became the repository for much of the detritus eroded from the new Alleghanian highlands. This repository most likely took the form of a vast, west-dipping, regional, alluvial-fan complex during Permian and Early Mesozoic time, nothing of which is now preserved (Faill, 1998). In fact, based on isotopic evidence, Dallmeyer (1989) concluded that the old orogen experienced major uplift during Late Permian to Early Triassic time and that by Late Triassic time (B220 Ma) parts of the old orogen in New England had been exhumed to nearly the present erosional level. Some of the resulting sediments may have been carried thousands of kilometers westward across the continental interior (Archer and Greb, 1995; Ettensohn, 2004), and Dickinson and Gehrels (2003) have shown that Appalachian sediments were transported as far as the western continental margin based on the ages of detrital zircons from Permian and Jurassic sandstones of the Colorado Plateau (see Blakey, Chapter 7). Eroded sediments were also transported eastward, as evidence indicates that rocks in southern New England had been buried by more than 5 km of sediments in Early Jurassic time (Roden-Tice and Wintsch, 2002). Moreover, in places on the American east coast, widespread, Early Mesozoic alluvial fans must have completely enveloped older Paleozoic rocks, and what is preserved today of these fans are mere remnants caught in synrift basins that were active as late as Cretaceous time (Roden-Tice and Wintsch, 2002). The breakup of Pangea in the central Atlantic region was part of major plate reorganization that had begun by Pennsylvanian time on the north and northeastern margins of former Laurussia and proceeded southward through Cretaceous time (Ziegler, 1989; Ettensohn, 1997). In the Atlantic region, breakup occurred both during the terminal phase of Pangean assembly and during the initial phases of Pangean dispersal along the tectonothermally thickened and highly elevated Alleghanian orogen (Manspeizer, 1994). Manspeizer (1994) has suggested that the breakup reflected continued northwest convergence in the lower crust concomitant with major uplift and southeasterly extension in the upper crust; as elevations increased, so did the vertical stress needed to drive orogenic collapse and extension. Extension may have also been abetted by the tectonic extrusion of former microcontinents or Grenville crustal fragments along dextral megashears (Manspeizer, 1994), which as they passed into the orogen seem to reflect reactivated shear zones that had previously accommodated Late Precambrian rifting and Paleozoic terrane accretion. As a result, about 40 offshore and onshore Late Triassic–Early Jurassic synrift basins formed on mostly accreted terranes in the old orogen along low-angle detachment surfaces that had been former Alleghanian thrust or strike-slip faults (Manspeizer, 1988, 1994; Manspeizer et al., 1989). Across most of the Atlantic area, basin sedimentation in half-graben basins began in the Late Triassic and ended by Middle Jurassic time when sea-floor spreading began (see McCracken, Chapter 14). The western onshore basins were generally high-relief, high-altitude, fluvial-lacustrine basins, whereas several of the offshore basins were apparently low-relief, sea-level, evaporite basins proximal to the future spreading center (Manspeizer, 1988; Manspeizer et al., 1989). The onshore basins contain a variety of terrestrial facies, including alluvial-fan and fluvial conglomerates and arkoses, sandstones and redbeds, as well as lacustrine mudstones, siltstones, and local coals arranged into simple transgressive-regressive cycles, probably driven by climate change (Van Houten, 1969, 1980). Many of the basins also exhibit Lower Jurassic flood basalts and related hypabyssal sills and dikes, which are parts of a larger period of spreading-related, post-orogenic, igneous activity that persisted from Early Triassic into Eocene time (Manspeizer et al., 1989). Onshore basin sediments, included in the Upper Triassic–Lower Jurassic Newark Supergroup, accumulated to thicknesses of 3–9 km and reflect more subhumid conditions in southern basins closer to the equator and more arid conditions in northern basins, although basin-related topography and later proximity to linear seas must have periodically forced unique climatic conditions (Manspeizer, 1994). Inasmuch as basin sediments overlie older Paleozoic crust and occur in post-orogenic, mostly intermontane basins developed on top of an inactive suture belt (Ingersoll and Busby, 1995), the Newark basins, in the broadest sense,

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are Alleghanian successor basins. Although orogen collapse and rifting had begun earlier across other parts of the Appalachian–Caledonian suture, along which the Iapetus-Rheic oceans had previously closed, Late Triassic rifting effectively ended the largely Paleozoic, Appalachian Wilson cycle and initiated the current Atlantic cycle. By Middle Jurassic time, spreading had generated an embryonic Atlantic Ocean, while in the Appalachian Basin, the regional Alleghany alluvial-fan was already undergoing deep dissection.

11. Economic Resources and Potential Resources abound throughout various parts of the Appalachian area, both sedimentary and crystalline (see Rose et al., 1968; de Witt and Milici, 1989; Feiss and Slack, 1989; Sanford, 1993c), but the very brief enumeration below will only deal with major mineral and hydrocarbon deposits present in foreland-basin parts of the area.

11.1. Energy resources 11.1.1. Coal Although thin coals are locally present in Devonian and Mississippian rocks, commercial Appalachian Basin coals are nearly all Pennsylvanian in age (Sholes and Skema, 1974; Miller, 1974). Coal was apparently first discovered from the Appalachian Basin in Pennsylvania in 1698, but major mining did not occur until the late 1700s (de Witt and Milici, 1989). Since 1890, more than 40.6 billion tons of coal have been produced from the Appalachian Basin, and typical annual coal production since 1991 has varied from 424 to 380 million tons, with West Virginia being the largest producer (EIA, 1995, 1999). The U.S. Department of Energy (1982) has indicated a demonstrated reserve base of about 116 billion tons from the Appalachian Basin, while EIA (1999) estimates are only slightly more than 108 billion tons, mostly bituminous coal, although other estimates suggest a reserve base nearly eight times that amount (Ferm and Muthig, 1982). Most of the coals come from three Appalachian subbasins: the Anthracite Basin of eastern Pennsylvania (Figure 3), the Pocahontas or Central Appalachian Basin, and the Dunkard or Northern Appalachian Basin (Figure 28). In the Anthracite Basin, anthracite has been mined for more than 200 years and occurs in four separated synclinoria. The anthracitization event was apparently Late Permian in age (Daniels et al., 1994) and has been ascribed to a formerly overlying ‘‘blanket’’ of rock or sediment (Levine, 1986) or to westwardly migrating hydrothermal fluids (Hower et al., 1993; Daniels et al., 1994). Faill (1998), however, has suggested that former diversion and thrusting of the crystalline externides over the coal basin may have contributed substantially to metamorphism. Unrelated, Upper Pennsylvanian anthracites also occur in the Narragansett Basin of Rhode Island and Massachusetts. The original coals here were deposited in alluvial settings in a strike-slip, pull-apart basin; metamorphism was apparently related to Alleghanian deformation and igneous activity (Murray et al., 2004). The Pocahontas, or Central Appalachian, Basin contains more than 30 major, Lower and Middle Pennsylvanian, bituminous coal beds. These coals developed in estuarine to fluvial settings and are organized into fourth- and fifth-order cycles controlled by tectonism and glacial eustasy (Chesnut, 1992, 1994; Greb et al., 2002, 2004; Figure 26). In eastern Kentucky alone, between 120 and 173 million tons of coal are produced annually from these beds, and many of them have relatively low ash and sulfur contents (Donaldson and Shumaker, 1981; Greb et al., 2002, 2004). In fact, about 24 percent of Appalachian coals are low-sulfur, 35 percent are mediumsulfur, and 41 percent are high-sulfur, with most of the low-sulfur coals coming from eastern Kentucky, Virginia, southern West Virginia, and the Pennsylvanian anthracite fields (EIA, 1999). Deposition of these coals was apparently mediated by eustatically controlled base-level rise and paleoclimate (Greb et al., 2004; Cecil and Dulong, 2003; Cecil et al., 2003, 2004). The coals thicken toward the center of the basin, where splitting into zones is common, but structure also exerts local control on thickness (Donaldson and Shumaker, 1981; Weisenfluh and Ferm, 1991). The Dunkard, or Northern Appalachian, Basin contains approximately 25 cyclic, Lower and Middle Pennsylvanian bituminous coals, which are correlative with coals in the Pocahontas Basin. However, about 20 cyclic, Upper Pennsylvanian coal beds and one Early Permian (?) bed, many of which are local or discontinuous, occur only in the Dunkard Basin (Frye, 1979). Although lower parts of the Upper Pennsylvanian section do show marine influence from the west or southwest, most of the Upper Pennsylvanian and Lower Permian coals are associated with alluvial and lacustrine environments (Kovach, 1979; Donaldson and Shumaker, 1981). In fact, the thickest and highest quality, Upper Pennsylvanian coals of the Monongahela Group were associated with starved-basin conditions in large fresh-water lakes and swamps (Kovach, 1979; Donaldson and Shumaker, 1981;

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Donaldson and Eble, 1991). The basin overlies the Rome trough, and accommodation space in the basin was probably controlled by reactivation of its bounding faults.

11.2. Coal-bed methane Coal-bed methane (CBM) is found in coals of Mississippian and Pennsylvanian age in the Appalachian Basin (Lyons, 1998; Markowski, 1998; Nolde and Spears, 1998; Attanasi, 1998), which is partitioned and assessed as two structural sub-basins, the Dunkard basin in the north and the Pocahontas or Central Appalachian Basin in the south (Figure 28). Methane degasification began as a safety measure in advance of underground mining in the Pocahontas Basin in the late 1970s and early 1980s and at first was merely vented into the atmosphere, but as the value of the resource was realized, CBM production from the Pocahontas Basin was begun from Virginia in 1988 and from West Virginia in 1995 (Nolde and Spears, 1998; Avary, 2004; Milici, 2004). As of 2002, 439 billion ft3 (BCF) of CBM had been produced from 2,359 wells in the Pocahontas Basin, and 8.5 BCF from 252 wells in the Dunkard Basin; most of that gas was biogenic methane, although thermogenic gas may be present in eastern parts of the basin where depth of burial, compaction, and coalification were sufficient (Milici, 2004). Presently, the gas-bearing units from these two basins are considered to be part of the Carboniferous Coal-bed Gas Total Petroleum System (Milici, 2004). Biogenic or microbial gas generation probably began as soon as initial peat deposits were formed and is probably continuing to some degree today; the coal beds serve as both source rock and reservoir. Trap formation began during deposition of the original peat deposits, and the seals are the connate waters that occupy the fractures and pore spaces in the coal beds as well as the shales that are intercalated with the coal; the critical stage in development of the system came during the Alleghanian orogeny when deformation created structures that enhanced fracture porosity in the coal beds (Milici, 2004). In the Pocahontas Basin, most CBM production is from the Lower Pennsylvanian Pocahontas and New River formations of the Pottsville Group in southern West Virginia and southwestern Virginia; estimated reserves include about 3.6 trillion ft3 (TCF) of undiscovered recoverable gas (Milici et al., 2003; Milici, 2004). In the Dunkard Basin, most of the CBM is produced from the Middle Pennsylvanian Allegheny Group and the Upper Pennsylvanian Monongahela Group in southwestern Pennsylvanian and northern West Virginia; reserves are estimated at 4.8 TCF of undiscovered, recoverable gas (Milici et al., 2003; Milici, 2004). Further exploration and assessment are ongoing in shelf areas of the Central Appalachian Basin in Tennessee, eastern Kentucky, and southern West Virginia and in Appalachian anthracite and semi-anthracite units from northeastern Pennsylvanian and westcentral Virginia, respectively (see Milici, 2004).

11.3. Hydrocarbons The purposeful production of natural gas in the Appalachian Basin was first begun in western New York in 1821, while the first production of oil in the basin occurred in western Pennsylvanian in 1859, although both substances had been accidentally encountered well before these times (de Witt and Milici, 1989). With nearly 150 years of production history, the Appalachian Basin is a mature oil- and gas-producing basin. Seventy-seven percent of the cumulative oil production occurred during the first 102 years from 1859 to 1960, with the peak in primary-oil production occurring in 1900 at 36,295,000 barrels (bbl), and the all-time peak of 37,585,000 bbl occurring in 1937 (Burns and Claus, 1985). Annual gas production peaked at 522 BCF in 1917. Cumulative oil production from the basin is now approaching 3.5 billion bbl, while cumulative gas production is now more than 41 TCF from more than 500,000 wells in 3100 fields (see Burns and Claus, 1985; de Witt and Milici, 1989). Petroleum source rocks comprise four major groups that are associated with major packages of siliciclastic strata, which are related to one or more orogenic events. Group 1 is associated with Pennsylvanian rocks, Group 2 with Lower Devonian to Mississippian shales, Group 3 with shales from Middle Ordovician–Upper Silurian siliciclastic and carbonate rocks, and Group 4 with various strata of Cambrian, and possibly Precambrian, age (Roen and Walker, 1996; Swezey, 2002). Swezey (2002) has also indicated two minor packages of siliciclastic rocks, the St. Peter and Oriskany sandstones, with likely sources from other groups. About 71 percent of the basin’s oil and 46 percent of its gas are derived from Devonian reservoirs, followed by Mississippian reservoirs, which have produced 15 percent of the oil and 32 percent of the gas production (Burns and Claus, 1985). Types of plays, however, are more varied and widely distributed than source rocks, but show stratigraphic proximity to same source-rock groups (Swezey, 2002). Many of the most typical plays are detailed in an atlas edited by Roen and Walker (1996). The presence of these distinct source-rock groups in the Appalachian Basin suggests that at least four different petroleum systems are present (Swezey, 2002). The presence of such systems in the basin is supported by the fact that oils below the Upper Silurian Salina evaporites are chemically different than oils from above them, suggesting that the evaporites form a regional hydrocarbon seal. Oils below the evaporites were apparently derived from

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pre-Salina source rocks, whereas oils above the evaporites came from post-Salina source rocks (Cole et al., 1987; Drozd and Cole, 1994). A similar relationship, however, may not hold for southern parts of the basin where Salina evaporites are absent (Swezey, 2002). Petroleum systems may be additionally subdivided into total petroleum systems (TPS) based on aspects of the source rocks, reservoir rocks, and traps. In the Appalachian Basin below the Salina, the Conasauga-Rome/ Conasauga, Sevier-Knox/Trenton, and Utica-Lower Paleozoic TPS have been defined; above the Salina, the Devonian Shale–Middle and Upper Paleozoic, Carboniferous coal-bed gas, and Pottsville (Alabama) coal-bed gas TPS have been defined (Milici et al., 2003). In terms of reserves, the U.S.G.S. (2003) estimated an average of 54 million bbl of undiscovered oil in the Appalachian Basin, about 86 percent of it in the Utica–Lower Paleozoic TPS and 14 percent in the Devonian Shale–Upper Paleozoic TPS. Similarly for gas, Milici et al. (2003) estimated an average of 70.2 TCF of undiscovered gas and 872 million bbl of undiscovered total natural gas liquids. About 94 percent of the undiscovered resource is continuous (unconventional) gas from the Devonian Shale–Middle and Upper Paleozoic, Utica–Lower Paleozoic, and Carboniferous coal-bed gas TPS (Milici et al., 2003). Recent gas discoveries in the Cambrian rocks of the Rome trough (Figure 14) have renewed interest in possible deep gas plays (Harris and Drahovzal, 1996; Harris and Baranoski, 1996) from other parts of the trough in the Appalachian Basin (Figure 1). Of the TPS mentioned above, gas shales from the Devonian Shale–Upper Paleozoic and Utica–Lower Paleozoic systems have been receiving increased attention. Gas from the Devonian–Mississippian black shales of the Appalachian Basin has been used since 1821, and the shales are considered to be world-class oil and gas shales (e.g., Dyni, 2005). Estimates suggest that between 577 and 1,131 TCF of natural gas are in place (Charpentier et al., 1993), but about only 34.1 TCF are probably recoverable (Milici et al., 2003; Milici and Swezey 2006; Sweezer, 2006). Controls on the origin of these shales are discussed by Ettensohn (1992). Work on the Utica shales is in a more preliminary state (e.g., Martin et al., 2003; Wallach and Rheault, 2003), but there is potential for 26.8 TCF of recoverable gas (Milici et al., 2003). If the ideas of Oliver (1986) are correct, occurrence of hydrocarbons in the Appalachian Basin may be related to presence in a basin whose margin has undergone tectonic deformation or uplift. According to Oliver (1986), fluids expelled tectonically or under the influence of hydraulic head generated by uplift and deformation of the basin margin may drive the migration of hydrocarbons into distal parts of the foreland basin and continental interior. Woodward (1958) has provided evidence for such westward migration, and the concentration of major hydrocarbon deposits along the western margin of the Appalachian Basin (Oliver, 1986, Figure 4) also supports the idea. In addition, the occurrence of oil and oil inclusions relative to likely paleotopography suggests at least two major periods of oil migration, one associated with the Taconian orogeny and a later period with the Alleghanian orogeny (Roedder, 1971; Haynes and Kesler, 1989), although oil migration in southern parts of the basin could just as easily be related to the Ouachita orogeny. Segmentation of the basin by reactivated basement structures not only facilitated the generation of some source and reservoir rocks (Donaldson and Shumaker, 1981), but no doubt also served to canalize or focus the fluid-flow regimes in the various basin petroleum systems.

11.4. Carbon sequestration Most of the energy used by humans is derived from the combustion of fossil fuels like those noted above, and the combustion releases carbon, mostly as carbon dioxide, into the atmosphere. Carbon dioxide, however, is a greenhouse gas that traps solar energy, thereby increasing the Earth’s surficial temperature and adversely affecting life on Earth. In the 200 years since the industrial revolution, the carbon dioxide content of the atmosphere has risen from 280 to 360 parts per million by volume, a 30 percent increase, causing great international concern and greater attention to carbon management (see Burruss and Brennan, 2003). One way of managing carbon is to store or sequester carbon dioxide in underground formations, which is called geologic sequestration, and the Appalachian Basin has a large potential capacity for geologic sequestration in the form of saline or brine reservoirs, depleted or abandoned oil and gas fields, coal beds too thin or too deep to be effectively mined, and carbonaceous (black) shales. In just Kentucky, Ohio, Pennsylvania, and West Virginia parts of the Appalachian Basin, a potential storage capacity of 209 gigatons of carbon dioxide is estimated, and this does not include the capacity of the nearly basin-wide carbonaceous shale deposits (MRCSP, 2005). Strategies are currently being developed, especially those that can simultaneously enhance oil, gas and coal-bed-methane recovery, and geologic field demonstrations are being planned in the Appalachian Basin (e.g., Wilson et al., 2003; MRCSP, 2005).

11.5. Mineral resources Mineral resources from the Appalachian Basin fall into four general types: evaporites, bedded ironstones, sandstone-hosted uranium-copper-vanadium, and carbonate-hosted lead-zinc-fluorite-barite. Major evaporite

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deposits occur in two stratigraphic settings, the Upper Silurian (Cayugan; Ludlow–Pridoli) Salina Group from New York to West Virginia, and the Middle Mississippian (Osage–Meramec transition; Lower Visean) Maccrady Formation in southwestern Virginia. In the Salina Group, halite and anhydrite were deposited in an elongate salt basin in the distal parts of the Salinic foreland basin stretching from western New York into western Pennsylvanian eastern Ohio and northern West Virginia (Alling and Briggs, 1961; Smosna and Patchen, 1978) (Figure 20). Salinic subsidence, combined with sediment starvation in a rain shadow, reef banks that periodically restricted the basin, and Late Silurian aridity, formed ideal conditions for evaporite deposition (Grabau, 1913; Ettensohn, 1994; Cecil et al., 2004). In contrast, gypsum, anhydrite, and halite in the Maccrady Formation are more localized and are thought to be related to sabkha and mud-flat deposition during lowstand conditions in an arid setting. Structural reactivation due to bulge moveout may have localized the deposition (Warne, 1990; Ettensohn, 1994; Ettensohn et al., 2002a), and later tectonic thickening created thicker, economic deposits (Feiss and Slack, 1989). Bedded ironstones have been sources of iron in the Appalachian region since 1765, and the most extensive deposits are the discontinuous, hematitic, ‘‘Clinton’’ (Early Silurian; Late Llandovery) iron ores that occur in carbonate and clastic deposits from New York to Alabama (Feiss and Slack, 1989). The origin of these deposits has been detailed elsewhere, but is probably related to the deep weathering of Taconian highlands in a seasonally wet climate and accumulation of the resulting weathering products in somewhat restricted coastal settings. Sideritic iron ores also occur throughout the basin, commonly along major unconformities like the Mississippian–Pennsylvanian boundary; at best they were marginal, low-grade ores and were locally exploited during the American Civil War. Sandstone-hosted ores of uranium-copper-vanadium occur in the Valley and Ridge province of eastern Pennsylvanian and western Virginia as roll-shaped lenses, strata-bound pods and lenses, and as mineralization associated with plant debris in channel sands from the Upper Devonian Catskill and Hampshire formations (Sevon et al., 1978; Feiss and Slack, 1989). According to Dennison and Wheeler (1975) the association of permeable fluvial sands, organic matter, and red- or brown-colored sands with greenish or grayish reduzate equivalents are important for uranium deposition from percolating ground water, and these are characteristics shared by parts of the Catskill and Hampshire formations. Based on these characteristics, the Pennsylvanian– Permian Dunkard Group, the Lower Pennsylvanian Lee or Pottsville sandstones, the Upper Mississippian Mauch Chunk–Pennington Group, and the Upper Devonian Hampshire–Catskill formations have the greatest potential for deposits of uranium and related elements. Deposits of carbonate-hosted lead-zinc-fluorite-barite abound in the Appalachian Basin, and many are effectively Mississippi-Valley-type deposits (Clark, 1987). Most of the ores occur in Upper Cambrian–Lower Ordovician, shallow-water, platform carbonates, commonly below regional unconformities. Some of the most important deposits occur in eastern Tennessee and western Virginia, and are associated with karstic breccias below the Knox unconformity (e.g., Hill and Wedow, 1971; Clark, 1987); other deposits occur along major facies transitions (Ohle, 1980). As with the migration of hydrocarbons, and perhaps associated with them (Roedder, 1971; Oliver, 1987; Haynes and Kesler, 1989; Davies and Smith, 2006), these deposits were probably associated with strongly saline brines that concentrated metals from basinal shales and then moved under tectonically established hydraulic gradients into favorable stratigraphic horizons in distal parts of the basin or in the continental interior (e.g., Oliver, 1986). Some of these brines were no doubt canalized into certain areas by major fault zones running across the basin (Heyl, 1972). Decreases in temperature, dilution, and chemical interaction with host rocks apparently led to ore precipitation in limited, stratigraphic habitats and as minor mineralization throughout the stratigraphic column (Roedder, 1971). Again, like the hydrocarbons, there is evidence of at least two phases of tectonically mediated deposition, one during Early–Middle Ordovician time (Hoagland, 1971; Carpenter et al., 1971; Hill et al., 1971; Harris, 1971) and another during Late Pennsylvanian–Early Permian time (Hearn et al., 1987).

12. Discussion and Summary The preserved Appalachian Basin extends nearly 1,675 km (1,030 mi) from northern Alabama in the U.S. to southeastern Ontario and southern Quebec in Canada, and may have extended another 1,000 km (600 mi) northeastwardly into Labrador, Canada, paralleling the Appalachian mountain belt. It is a composite, retroarc foreland basin, which formed in stages over a period of about 340 Ma during the Iapetan or Appalachian Wilson cycle from latest Neoproterozoic (B570 Ma) to Late Triassic (B230 Ma) time. The basin is one of the best studied foreland basins in the world and is part of the ‘‘type area’’ for geosyncline theory and Wilson cycles. In addition, it has served as a natural laboratory for the evolution of ideas concerning orogenies, isostasy, structural mechanisms, flexural modeling, and the migration of basin fluids.

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Sedimentary components of the basin developed on Mesoproterozoic, Grenvillian crust and did not begin developing until about 600 Ma when Late Neoproterozoic–Early Cambrian rifting formed the Iapetan (Appalachian) passive margin. This rifting generated an irregular, zigzagged, passive margin, as well as a series of basement faults and zones of crustal weakness that would not only influence the course of later Paleozoic convergence events (orogenies) on the margin, but also later basin configuration and sedimentation, basin fluid migration, and even later collapse of the Alleghanian orogen to form the current Atlantic passive margin. Basin development and sedimentation are both closely related to four Appalachian-margin convergence events through deformational loading that generated accommodation space and through deformational uplift that in large part generated sediment infill. The nature and distribution of the accommodation space and sediment infill, however, were both strongly mediated by the pre-existing structural framework, Paleozoic paleoclimate and paleogeography, and Paleozoic sea-level variations; in some cases, there is clear evidence of reciprocal interactions between mediating conditions, orogenic events, and basin sedimentation. Sedimentation history in and near the foreland basin generally reflects four phases of development: rifting, passive margin, convergent margin, and orogenic collapse. The first phase represents the rifting away of Gondwanan parts of Rodinia during collapse of the Grenvillian orogen at the end of the Laurentian Wilson cycle; the rifting not only contributed to the formation of Laurentia, but also to evolution of the Iapetus Ocean on its eastern margin and the subsequent formation of island arcs on both sides of the new ocean basin during the nearly 300 Ma period of rifting and passive-margin development. Rifting lasted for about 230 Ma from Late Neoproterozoic to Early Cambrian (765–535 Ma) time and included the clastic infill of continent-interior and continent-margin rift basins (Figure 14). Overlapping with rift development, the passive-margin phase lasted for about 100 Ma, from latest Neoproterozoic time to the Early–Middle Ordovician transition (B570–472 Ma) and saw the development of a stable platform with uppermost Neoproterozoic to Middle Cambrian, terrestrial and shallow-marine, clastic sediments overlain by shallow-water, Upper Cambrian to Lower Ordovician carbonates (Figure 14); in a few situations, deeper water, continental-margin equivalents are also preserved. The Late Cambrian change to carbonate deposition, which would predominate across large parts of the Appalachian Basin area until latest Mississippian time, reflects rising sea level and movement of the Appalachian region of Laurentia/ Laurussia into the generally evaporative subtropical, trade-wind belt, where it would more or less remain until latest Mississippian time. Rifting and subsequent evolution of the Iapetus Ocean during a period of nearly 300 Ma set the stage for subduction and island-arc development on southeastern Laurentian and northern Gondwanan margins of the ocean basin. On the western or Laurentian side of the ocean, subduction and island-arc formation had probably begun by Late Cambrian time (Penobscottian orogeny?), and on the Gondwanan side, even earlier. However, rotation of Gondwana, rifting along its northern margin, and the possible inboard shifting of transforms transferred the northern, peri-Gondwanan-arc, or Avalonian, terranes to an oceanic plate and sent them on their way toward Laurentia during Middle Cambrian to Middle Ordovician time (Nance et al., 2004). The southernmost of these arc-terranes, the Carolina terrane, was probably approaching the Laurentia margin from the south by Late Ordovician time (Hibbard, 2000; Hibbard et al., 2002), whereas other Avalonian terranes seem to have arrived from the north by late Early Silurian time. Closure of the Iapetus Ocean, or the ocean basin between these terranes and the Laurentian margin, was effectively complete by Early Devonian time. However, the ocean basin on the other side of the terranes, separating them from Gondwana and known as the Rheic Ocean, did not close until Pennsylvanian time. Thus, in generating the Iapetan and Rheic ocean basins, as well as subsequent arcs along their margins, the foundation was set for several episodes of convergence that would generate the Appalachian foreland basin. Oceanic convergence with some of the outboard microcontinents or Laurentian continental fragments stranded in oceanic crust after Iapetan rifting probably began as early as Middle Cambrian time (van Staal, 2005), but the earliest collision involving the Laurentian margin and nearby arcs in the western Iapetus did not begin until near the Early–Middle Ordovician transition (472 Ma). At this time, in response to arc obduction onto the Cambro-Ordovician, carbonate-platform margin, carbonate facies that had predominated on the eastern margin of Laurentia were uplifted from the east, exposed briefly, and suddenly dropped to depths too deep to support carbonate deposition. This event marks inception of Appalachian (Iapetan) convergence, and with it, the Laurentian platform margin and adjacent parts of the craton differentiated for the first time into a mosaic of foreland-basin and cratonic clastic and carbonate facies, the distributions of which were largely controlled by old basement structures reactivated by the orogeny; the kind of uniform, clastic, and carbonate deposition that characterized postrift, passive-margin Laurentia would not again characterize the Appalachian area again until post-Paleozoic time. This change from a passive-margin to a convergent tectonic regime marks the true beginning of the Appalachian foreland basin and is reflected nearly basin-wide in a major unconformity variously called the St. George, post-Sauk, Owl Creek, sub-Tippecanoe, or Knox unconformity (Figures 4, 14, and 16).

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Appalachian convergence occurred during two, second-order, orogenic cycles of approximately 50–60 Ma in length. The earlier Caledonian cycle (Middle Ordovician–Early Devonian; 472–411 Ma) reflects closure of the Iapetus Ocean through subduction involving island-arc complexes and peri-Gondwanan, Avalonian terranes, whereas the later Variscan–Hercynian cycle (Early Devonian–Late Permian?; 411–260 Ma) reflects closure of the Rheic Ocean through collision with Gondwana as well as collision-related transpression and strike-slip adjustment of previously accreted Avalonian terranes. The larger second-order cycles, moreover, can be divided into third-order orogenic cycles, reflecting individual collisional phases within larger orogenies. The Caledonian cycle contains the Taconian and Salinic orogenic events, and both probably reflect some aspect of sinistral transpression involving accretion of peri-Gondwanan terranes. The Salinic event is in effect a more southerly manifestation of the larger Caledonian orogeny to the north; it ended with the accretion of Baltica (Western Europe) to northern Laurentia and the accretion of several peri-Gondwanan, Avalonian terranes to the southern, Appalachian margin of Laurentia, together forming a larger continent called Laurussia. A possible third event that I have informally called the ‘‘Helderberg’’ event probably represents collapse of the Taconian and Salinic orogens and development of successor basins throughout the old orogen due to slab separation and delamination (Figure 21A). The Helderberg event probably represents a necessary transitional period in the change of tectonic regimes. The new Variscan–Hercynian cycle contains the Acadian (Neo-Acadian) and Alleghanian events, both of which reflect major components of dextral strike-slip and transpression, most likely related to the final docking and rotation of Gondwana toward Laurussia. Although orogenies can be distinguished to some extent in the crystalline Appalachians based on structural analysis, geochemistry, and radiometric dating, flexural stratigraphic analysis, which indicates the occurrence of convergence in smaller, discrete phases that leave distinctive sedimentary sequences in the foreland basin, enables more detailed discrimination of tectophases and orogenies from the foreland basin. In fact, mapping changing sedimentary-sequence distribution in time and space clearly supports the idea of diachronous, Appalachian orogenies that migrated along strike due to oblique convergence and transpression. Flexural modeling from the basin suggests that Appalachian orogenies occurred in a series of two to four tectophases that were mediated by continental promontories, relics of Iapetan rifting, which focused convergence and restricted its along-strike continuation. Appalachian tectophases ranged in duration from about 2 to 42 Ma, with an average of about 12 Ma, and the cyclic, sedimentary manifestation of 13 such tectophases during 4 orogenies largely comprise the infill of the Appalachian Basin, which we see today (Figures 4 and 14). Modeling suggests that the foreland basin responded to cycles of flexural loading and relaxation during any complete tectophase with a seven-part sequence of lithologies, which is bound by regional unconformities and manifest in the foreland basin as a generalized, third-order, transgressive-regressive cycle. These cycles have been called foreland-basin tectophase cycles or sequences (Ettensohn, 1994, 2005). Hence, the basin is filled with 13, stacked, third-order tectophase cycles, arranged into 2 larger, second-order, orogenic cycles (Figure 4). Not every tectophase cycle is complete, but 11 cycles from the first 3 orogenies exhibit at least lower parts of the typical, ascending, cycle manifestation: a basal dark marine shale, a flysch-like clastic unit, and a molasse-like unit of marginal-marine to terrestrial clastic sediments with redbeds. The dark shales represent the time of greatest deformation and loading and are the most distinctive unit in the cycle; hence, mapping their distribution in space and time shows basin migration and the along-strike progression of tectophases during orogeny. The 11 earlier Appalachian cycles are typical of subduction-type orogenies, in which deformational loads must first mount a continental-margin ramp, and consequently accumulate across short distances at the margin, especially at promontories. Loads assembled largely at the margin, however, generate deeper, isostatically compensated foreland basins, in which sediment-starved conditions that favor dark-shale deposition rapidly develop. Such early orogenies are typically ‘‘symmetrical’’ with sediment sinks in the subduction zone on one side and in the foreland basin on the other. In the Appalachian situation, moreover, foreland basins that developed during the first four orogenies formed nearly perpendicular to climatic zones, largely in the southern, arid, subtropical belt that lacked year-round precipitation (Scotese, 2003). This condition militated against maximum clastic influx to Early Appalachian foreland basins such that the basins were rarely overfilled, carbonates were relatively common, sediment-starved conditions periodically developed, and except during a few lowstands, marine conditions almost always prevailed. This scenario changed during the last two Alleghanian tectophase cycles, when throughout the Appalachian Basin, molasse-like, marginal-marine to terrestrial clastic sediments came to dominate, carbonates and other marine sediments became uncommon, and a pronounced fourth- to fifth-order cyclicity overprinted each tectophase cycle. Although similar flexural events seem to have generated the Alleghanian tectophase cycles, the above differences apparently reflect the late-stage, collisional nature of the orogeny, as well as both the tectonic and climatic ‘‘asymmetry’’ of the orogen. In addition, the prominence of strike-slip, escape tectonism in the orogen may have also been significant. Nevertheless, by the time of Alleghanian collision, the deformational load had almost certainly surmounted the marginal ramp and migrated far enough across the adjacent platform that the

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more dispersed load could only generate broad, shallow foreland basins incapable of maintaining deeper water, marine regimes. In addition, continental collision and resulting highlands to the south and east must have prevented the symmetrical distribution of clastic debris to the ocean or subduction side of the orogen, with the result that most of this debris would have found its way via northern and transverse drainage into the foreland basin. Moreover, because of continental rotation during collision, the foreland basin and orogen became oriented nearly parallel to the equator in the rainy, humid, tropical, equatorial belt such that weathering and transport of clastic debris from the orogen into the foreland basin must have been far greater than during early orogenies. As a result, Alleghanian foreland basins were shallow and generally close to overfilling, or overfilled, with terrestrial to marginal-marine clastic debris throughout much of their history. Overall, Appalachian Basin infill as well as the character and composition of its second- and third-order, sedimentary cycles were largely controlled by orogeny type, timing of the orogeny within the larger Wilson cycle, and disposition relative to climate belts. However, lower level, fourth- to sixth-order cycles were probably controlled by tectonism (periods of thrusting) and eustasy. Although Alleghanian tectonism probably continued through Late Permian time, especially in northern parts of the orogen, much of the orogen had become tectonothermally thickened and uplifted, generating enough vertical stress to cause orogen collapse and extension, especially in the crystalline internides. Collapse may have also been aided by microplate extrusion along dextral megashears, and much of the collapse and extension apparently occurred along older thrust faults, terrane-accretion boundaries, and even Iapetan rifts. The Alleghanian foreland basin, as well as the newly established zone of Atlantic rifting must have been awash in extensive, thick, alluvial-fan deposits that spilled over onto the continental interior, but in the Appalachian Basin, none of these deposits have been preserved. By Late Triassic time, the Iapetan or Appalachian Wilson cycle had ended, and the Atlantic cycle had begun. Except for orogen collapse, evidence for all other phases of the Iapetan or Appalachian Wilson cycle are preserved in the Appalachian foreland basin, making basin events, stage by stage, an ideal model for foreland-basin development through time. In addition, the basin reflects the larger events of the Appalachian area itself and provides a model for continent-margin response to the progression of a Wilson cycle. Hence, the broad patterns of basin development and the continent-margin events that generated them are outlined above based on current knowledge and interpretations. It is important to note, however, that at nearly every step in basin development, the outline of the old Iapetan margin and old basement structures left over from the end of the Laurentian (Grenvillian) Wilson cycle contributed immensely to the evolution of the new Paleozoic basin. Although regional, and perhaps global, tectonic events were prime drivers for Appalachian Basin development, the pattern that emerges from the basin thus far also reflects the critical imprints of paleoclimate, eustatic fluctuation, and local structure. As the early workers soon came to recognize, the Appalachian area is an ideal natural laboratory in which to understand these interactions; as problems continue to exist at all scales, an even fuller understanding will only come with more work.

ACKNOWLEDGMENTS I wish to thank D. Lavoie and A.D. Miall, who reviewed an earlier version of the manuscript and suggested many useful improvements.

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CHAPTER 5

The Paleozoic Western Craton Margin Andrew D. Miall

Contents 1. Introduction 2. Historical Background 3. The Rifted Margin of Laurentia 3.1. Southern Canadian Rocky Mountains and Great Basin 3.2. Yukon Territory and northwest territories 4. The Sauk Sequence and the Cambrian–Ordovician Shelf-to-Basin Transition 4.1. The Kicking Horse Rim and Burgess Shale of the southern Canadian Rocky Mountains 4.2. Northern British Columbia 4.3. Yukon and northwest territories 4.4. Great Basin: Nevada, Utah, Idaho 5. Middle Ordovician–Early Devonian (Tippecanoe Sequence) 5.1. Northern Canada 5.2. Great Basin 6. Grand Cycles 7. Lower to Upper Devonian (Kaskaskia-I Sequence) 7.1. Northern Canada 7.2. Peace River Arch 7.3. Ancestral Uinta uplift 7.4. Great Basin 8. Mississippian Arc Collisions and Termination of Parts of the ‘‘Passive’’ Laurentian Margin (Kaskaskia-II Sequence) 8.1. Great Basin 8.2. Western Canada 9. Pennsylvanian–Permian (Absaroka-I and II Sequences) 10. Triassic–Jurassic: Termination of the ‘‘Passive’’ Continental Margin 11. Conclusions Acknowledgments References

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Abstract The western margin of the Paleozoic Laurentian continental margin is now largely buried beneath the accreted terranes and fold-thrust belt of the Cordilleran Orogen. An extensional continental margin developed from Mexico to Yukon during the breakup of Rodinia about 900–800 Ma, commencing with a series of Precambrian rift basins, followed by the development of a wedge of Paleozoic continental margin (‘‘miogeoclinal’’) sediments up to 6 km thick. Shallow-water carbonate and clastic sediments characterize most of the preserved sediment pile. Slope and deep-basin sedimentary rocks are rarely preserved. An arc collided with this continental margin from southeastern California to central Idaho between the latest Devonian or Early Mississippian generating the Antler Orogeny and the emplacement of the Roberts Mountains Allochthon on the continental margin, above a thrust belt. Most of the Canadian portion of the margin remained an extensional margin until the Nicola Arc was thrust eastward over the margin during the Jurassic. The Canadian Margin may have developed by simple-shear extension during the Paleozoic, that portion lying south of the 60th parallel forming the upper-plate margin and the margin north of the 60th parallel constituting the lower-plate margin, with the Liard Line functioning as a transfer fault across which the shear polarity reversed direction. Sedimentary Basins of the World, Volume 5 ISSN 1874-5997, DOI 10.1016/S1874-5997(08)00005-1

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1. Introduction The western continental margin is particularly well exposed and has been extensively studied in four main areas (Figure 1), for which the stratigraphy is summarized in Figure 2: 1) The Great Basin, which includes most of Nevada, and extends from the eastern limit of the Basin-and-Range tectonic province, the Wasatch Line, east of Salt Lake City, Utah, to the eastern margins of the Sierra Nevada, California. The rocks of the Paleozoic continental margin of Laurentia are particularly well exposed in the tilt blocks of the Basin and Range within the Great Basin of Nevada, plus adjacent areas of Utah and Idaho. A symposium on Paleozoic paleogeography edited by Stewart and Suczek (1977) and the Cordilleran volume of the Decade of North American Geology project (DNAG) (Burchﬁel et al., 1992) provide basic reference sources for this area. It should be noted that the eastern limit of Basin-and-Range extension was determined by the western margin of unthinned continental crust, and the Great Basin was named for a physiographic area within the Basin-and-Range tectonic province, that is characterized by internal drainage. 2) The Rocky Mountains of Banff, Jasper and Yoho national parks, Canada. Useful syntheses of the Western Canada Sedimentary Basin, which includes much information on the ancient continental margin, were compiled by Ricketts (1989) and by Mossop and Shetsen (1994). The DNAG volumes on the Canadian Cordillera (Gabrielse and Yorath, 1992) and the Sedimentary Cover of the Craton (Stott and Aitken, 1993) are also essential sources. 3) The Mackenzie Mountains of the Northwest Territories, north of the 60th parallel, for which the DNAG volumes are the essential references. The regional tectonic history of this area is discussed in detail by Cecile et al. (1997). 4) The mountainous areas of northern Yukon and adjacent areas of the Northwest Territories. Again, the DNAG volumes plus Cecile et al. (1997) are the most useful sources. This region includes large areas of ‘‘pericratonic terranes’’ (Figure 1), that is, terranes that are interpreted as originally part of Laurentia, but which have been displaced at some time during the Phanerozoic.

2. Historical Background The application of plate-tectonic concepts to the interpretation of the ancient geological record, beginning in the late 1960s, has provided many signiﬁcant breakthroughs in our understanding of complex geologic regions. The western margin of the North American continent was amongst the earliest tectonic provinces to be understood using these new ideas. Stewart (1972) and Burchﬁel and Davis (1972) suggested that, during the Late Precambrian and Early Paleozoic, the margin of the ancient continent extended through the central portion of the Cordilleran Orogen (Figure 1). Interpretation the rocks of coastal belts extending from Yukon Territory to California, had to await the development of terrane concepts in the late 1970s (Monger and Price, 1979; see Chapters 10 and 11). Belts of Upper Precambrian and Lower Paleozoic strata, reaching thousands of meters in thickness extend NNW–SSE through the foreland region of the Rocky Mountains from the Yukon to Idaho, and on southwesterly through central Nevada (Figure 3). Stratigraphic and facies patterns of these rocks suggest an interpretation in terms of a ‘‘passive,’’ extensional continental margin. Stewart (1972) and Stewart and Suczek (1977) made an explicit comparison to the modern Atlantic continental margin of eastern North America (see Miall et al., this volume, Chapter 14), including (1) evidence of rift faulting and subsidence in the Late Precambrian, which they compared to the Triassic ‘‘Newark’’ rift basins of New England and Atlantic Canada; (2) fault-bounded basins containing great thicknesses of westerly derived Upper Precambrian clastic strata, including several intervals of glacially derived diamictites; (3) basaltic intrusions and lava ﬂows of Late Precambrian age, which can be compared to similar Triassic–Jurassic igneous activity in the Newark basins; (4) gradual onlap of Lower Paleozoic strata onto the continental margin during the thermal subsidence phase of margin development; (5) Lower Paleozoic shallow-water strata of ‘‘miogeoclinal’’ character, consistently thickening from the craton as they are traced westward into the deformed foreland belt, similar to the pattern of modern sedimentation on the Atlantic continental shelf off the United States; and (6) a transition, where preserved, into deeper-water, largely clastic facies toward the west, suggesting a comparison with continental slope and basinal facies of the deeper parts of the modern Atlantic Margin. Bond and Kominz (1984) and Bond et al. (1985) applied the methods of backstripping to the ancient continental margin and conﬁrmed the appropriateness of the model of crustal stretching and thermal subsidence. As discussed later, the development of the margin is now known to include more than one rifting event, and the probable presence of crustal detachment faults. We now know that the rocks to the west of the Early Paleozoic

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Figure 1

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Location map.

continental margin consist of a belt of amalgamated terranes as much as 500 km wide, attesting to the accretionary growth of the continent since the Early Paleozoic. From the Late Precambrian until the Late Devonian, the entire western margin functioned as a passive or extensional margin (Figures 5, 7 and 9 in Chapter 1). Then, commencing in the latest Devonian or Mississippian, the Antler Arc began to collide with the continental margin in the region of what is now central Nevada (Figure 11 in Chapter 1), emplacing the Roberts Mountains

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Figure 2 Correlation table, showing selected stratigraphic names. Sources: Time scale from www.stratigraphy.org; Sloss Sequences from Sloss (1988), with ages revised according to www.stratigraphy.org. Nevada, Miller et al. (1992); Alberta and British Columbia, Bond and Kominz (1984), Ricketts (1989); Northwest Territories,Yukon, Gabrielse and Yorath (1992).

Allochthon above the miogeoclinal margin (Figures 1 and 3; Figure 11 in Chapter 11; see this volume, Chapter 11). In Nevada, this was followed by extensional and transcurrent deformation associated with the ‘‘Phase two’’ tectonism that primarily affected the southern part of the continent (Chapters 7 and 8), and then the Permo-Triassic Sonoma Orogeny. Meanwhile, the evidence suggests that most of the remainder of the western continental margin, extending from Oregon to Yukon, remained in an extensional regime until the Early Jurassic (Gabrielse and Yorath, 1992). A west-facing arc (the Nicola or Quesnellia Arc) had developed off the coast from Oregon to central British Columbia by the Late Paleozoic or Triassic (Figure 21B in Chapter 1), but until the Early Jurassic this appears to have been an extensional arc, with a backarc basin (Slide Mountain terrane) located to the east. The continental

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Figure 3 (1972).

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Isopach of Upper Precambrian to Lower Cambrian strata on the western continental margin. Redrawn from Stewart

margin of a backarc basin functions tectonically and stratigraphically much like a passive continental margin. Extensional tectonism of the continental margin associated with this plate-tectonic regime included subsidence of the Liard Basin, Prophet Trough and the Peace River Embayment. Collision and subduction-related basins of the North American Cordillera are described in Chapters 10 and 11.

3. The Rifted Margin of Laurentia 3.1. Southern Canadian Rocky Mountains and Great Basin As noted earlier, the modern Atlantic Margin of the United States was used as an analogue for the ﬁrst platetectonic interpretation of the Precambrian–Lower Paleozoic margin of Laurentia. Subsurface data from the

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Atlantic margin also provided the basis for the ﬁrst quantitative models for subsidence and sedimentation on passive or extensional, ‘‘Atlantic-type’’ continental margins (Watts, 1981; Watts et al., 1982). The procedure, which is now standard, by which rates of subsidence and sedimentation are calculated, and related to tectonic and isostatic controls, was ﬁrst developed using these data (Steckler and Watts, 1978). This procedure, called ‘‘backstripping,’’ was applied to an investigation of the evolution of the ancient continental margin of British Columbia by Bond and Kominz (1984), in order to explore the appropriateness of a Watts-type Atlantic-Margin model for these rocks. Stewart and Suczek (1977) constructed a subsidence curve for the continental margin of southern Nevada, and also interpreted the result in terms of an Atlantic-Margin-type model. Bond and Kominz (1984) introduced one important modiﬁcation to the backstripping method, which was required by the major difference in rock type between the Jurassic to modern sediments of the Atlantic Margin, and the Precambrian–Cambrian rocks of British Columbia. On the Atlantic Margin, the succession is primarily clastic. The sediments there have high initial porosities, which are gradually reduced by burial compaction. This compaction results in thickness reductions during burial, which can be compensated for in the subsidence calculations by using the local empirical relationship between porosity and depth. However, most of the miogeoclinal succession on the ancient British Columbia Margin consists of carbonate sediments which typically undergo cementation and lithiﬁcation early. Bond and Kominz (1984) constructed a different porosity-depth curve for these rocks based on studies of early carbonate diagenesis and tested against data from a well drilled off the Florida Margin through a predominantly carbonate section. This allowed them to ‘‘delithify’’ the strata as a ﬁrst step in the backstripping procedure. The oldest sedimentary rocks within the miogeoclinal belt are the Belt–Purcell Supergroup of Helikian Age (1.8–1.5 Ga). These constitute an immense thickness of largely clastic deposits up to 15 km thick. They were formed in rift and epicontinental basins within the Rodinia Supercontinent, and underwent a deformational event about 1,300 Ma. These rocks predate the Grenville Orogeny (1.3–1.0 Ga), which was the ﬁnal event in the assembly of Rodinia, and are therefore unrelated to the formation of the western Laurentian Margin. The Belt– Purcell Supergroup is overlain with angular unconformity by the Windermere Supergroup, of Helikian (800– 575 Ma) Age and up to 9 km in thickness (Gabrielse, 1972). Extensive maﬁc intrusives and extrusives are related to deep-crustal fractures that are thought to have been active during sedimentation. Approximately 6 km of Lower Paleozoic miogeoclinal strata lie above these rocks and are spectacularly exposed in the Rocky Mountains of Banff, Jasper and Yoho national parks. A succession of shallow-water, mainly carbonate sediments, extends from the craton westward to near the Alberta–British Columbia border, where a major facies change takes place, over what Aitken (1971) termed the ‘‘Kicking Horse Rim’’ (named after the famous mountain pass of the same name). The Paleozoic section thickens and changes facies westward into deeper-water facies, the details of which are described in a later section. Rocks of Cambrian to earliest Silurian Age are well represented, corresponding to the Sauk I to Kaskasia II sequences. The major sequence-bounding unconformities of Sloss (1963) can be recognized within the succession. A major angular unconformity occurs near the middle of this succession; rocks of the Tippecanoe II Sequence, corresponding to most of the Silurian System, are largely absent from the southern Canadian Rocky Mountains. Bond and Kominz (1984) constructed a restored stratigraphic cross-section through the southern Rocky Mountains, approximately along the transect followed by the Trans-Canada Highway. Figure 4 shows the Precambrian to lowest Silurian portion of this cross-section (above which there is a major unconformity). Their subsidence curves suggest that this interval corresponds to the main thermal subsidence phase of an extensional margin. Rapid thermal subsidence appears to have commenced at some time between the latest Precambrian and the Mid-Early Cambrian (600–550 ma). A thin unit, the Hamill Group, which spans this age range and occurs in the southern Rocky Mountains of British Columbia, is cited by Bond and Kominz (1984) as providing support for this interpretation. It consists of ‘‘coarse-grained arkosic sediments and scattered maﬁc lavas, an association that is suggestive of a continental terrane undergoing rifting’’ (Bond and Kominz, 1984, p. 167). The development of oceanic crust corresponding to the initiation of Panthalassa, would have followed shortly after the deposition of this unit, initiating the thermal subsidence of the adjacent margin. In British Columbia, the increase in thickness of the Lower Paleozoic section from the miogeoclinal carbonate bank to the deep-water shale basin may indicate a greater rate of subsidence of thinned continental crust at the outer margin of Laurentia. In this interpretation, the Kicking Horse Rim is located at approximately the boundary between the normal and thinned basement created by the crustal extension that commenced in the Late Precambrian. Stewart and Suczek (1977) had earlier developed a similar interpretation for the corresponding continental margin of Nevada. They constructed a subsidence plot (Figure 5) which they interpreted as indicating an exponential patterns of subsidence. A succession of terrigenous sediments of Late Precambrian Age extending from Washington to southern California (Figure 6) is interpreted as the product of rifting and, subsequently, of erosion of the thermal uplift that typically precedes continental separation. They interpreted the beginning of rifting as taking place at about 900–800 Ma, with the commencement of post-rift thermal subsidence starting

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Figure 4 Reconstructed stratigraphic cross-section through the Precambrian--Lower Paleozoic section of southern Alberta and British Columbia. This section represents a palinspastic reconstruction. Locations of two of the major thrust sheets are indicated for reference purposes. The MacConnell Thrust is the easternmost of the thrusts to bring Paleozoic strata to the surface, and de¢nes the front of the physiographic Rocky Mountains west of Calgary. Redrawn from Bond and Kominz (1984).

Figure 5 Subsidence plot for the Uppermost Precambrian to Devonian strata of the western United States (redrawn from Stewart and Suczek, 1977). At the time this plot was constructed, dating of the oldest rocks in the succession was uncertain, and two positions are shown for the commencement of subsidence.

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Figure 6 The extent of pre-drift sediments on the western margin of Laurentia in the United States (stippled area) and the onlap pattern of Cambrian strata. Adapted from Stewart and Suczek (1977).

Figure 7 Simple-shear model of extensional continental margin development, after Lister et al. (1991) and Cecile et al. (1997). The British Columbia--Alberta portion of the continental margin may have been an upper-plate margin, whereas the area north of 601N may have functioned as a lower-plate margin, the reversal in facing direction taking place at a transfer fault corresponding to the Liard Line (Cecile et al., 1997).

near the end of the Precambrian. Watts (1981, 1989) and Watts et al. (1982) showed that as the ﬂexural rigidity of the continental crust increases and the outboard sediment load increases following continental separation and drift of the continent away from the seaﬂoor spreading center, the ﬂexural hinge at the edge of the continental crust migrates gradually cratonward, resulting in gradual onlap. This is clearly shown in the map of the western United States (Figure 6) and in the cross-section of western Canada (Figure 4) constructed by Bond and Kominz (1984). In a subsequent study, Bond et al. (1985) extended their analysis to a series of locations between Yukon and Utah, and conﬁrmed the general pattern of thermal subsidence commencing between about 600 and 555 Ma,

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indicating the widespread initiation of a passive margin along the western margin of Laurentia. The general similarity of the stratigraphy along the length of the western continental margin as far north as 601N suggests that the rifting and thermal subsidence model applies to the entire western Laurentian Margin, although Cecile et al. (1997) argued that the pure-shear model of Bond and Kominz (1984) is not appropriate for the northern Cordillera, as summarized later. The length of time between the deposition of the oldest of the presumed rift-related Windermere Supergroup (800 Ma) and the initiation of seaﬂoor spreading (660–550 Ma) is not a problem. Although early models for extensional margins considered rifting as if it was ‘‘instantaneous,’’ for the purposes of calculation (e.g., McKenzie, 1978), in fact, it is not at all uncommon for the pre-drift rift stage to extend for as long as 200 Myr, and to include several discrete rifting episodes, as has been documented for the margins of the North Atlantic, between Newfoundland, Spain, Greenland and the European Margin west of Britain (e.g., Surlyk et al., 1981). Cecile et al. (1997) offered a similar argument for the length and complexity of the extensional tectonism of the northern Cordilleran Margin. Deep-crustal reﬂection proﬁling, of the type pioneered by Oliver (1982), led Lister et al. (1991) to recognize that a range of tectonic mechanisms occurs at extensional continental margins. To the original pure-shear model of McKenzie (1978) has been added a range of simple-shear models based on the recognition that the rifting and separation of some continents takes place across master detachment faults that penetrate the entire crust at a high angle. In the absence of seismic data (which are not available for the study area) structural trends and isopach patterns may be quite distinctive. Many characteristics of the Paleozoic continental margin of British Columbia and the territories to the north indicate that crustal separation along this portion of the margin may have followed a pattern of simple-shear evolution. The western edge of the Alberta–British Columbia portion of the margin is characterized by zones of uplift, including the West Alberta Arch and the Macdonald Platform that, during the Cambrian and Ordovician, functioned as positive elements over which developed widespread Sub-Devonian unconformities. These and other indicators, discussed by Cecile et al. (1997), suggest that the margin in this region may have functioned as the upper-plate margin above an east-dipping detachment fault (Figure 7). Lithoprobe data from this area do not throw much light on the structure of the Paleozoic continental margin, owing to postPaleozoic deformation, but a reconstruction of the Paleozoic Margin derived from SNorCLE line data offered by Welford et al. (2001, Figure 13) is consistent with this interpretation. Aeromagnetic studies have indicated the probability of mantle underplating and crustal thickening in northeastern British Columbia (Saltus and Hudson, 2007), exactly where a simple-shear model of upper-plate architecture would predict (Lister et al., 1986). Structural and isopach trends suggest that Paleozoic sedimentation and tectonism were inﬂuenced by a series of northeast–southwest-trending lineaments that probably originated as faults or terrane boundaries in the Precambrian basement. The most important of these, which has been named the Liard Line, crosses the continental margin at about the latitude of the British Columbia–NWT border (Figure 1). As discussed later, this line may have functioned as a major ‘‘transfer fault’’ during the Early Paleozoic. Another element of the Bond and Kominz (1984) model is also worth commenting on. They compared a modeled two-dimensional cross-section of the continental margin with an actual, restored cross-section (Figure 8). Note that the restored section exhibits a much greater than predicted thickness of the marginal ﬂexural wedge. This probably indicates that both thermal subsidence and crustal thinning were underestimated in the backstripping calculations. The thin wedge of Middle to Upper Cambrian strata extending across the craton is probably the result of the episode (or episodes) of eustatic sea-level rise (or dynamic-topography subsidence) that were responsible for the development of the Sauk II and Sauk III sequences. Following this episode of high sea level, there was a widespread regression. Rocks of Middle Ordovician to Early Devonian Age are almost entirely absent from western and central Alberta, owing to long-lived uplift, or to repeated episodes of sedimentation and erosion of the West Alberta Arch (or Ridge). They are, however, widespread north of the 60th parallel.

3.2. Yukon Territory and northwest territories Yukon Territory includes a large area of the continental margin that has been interpreted as a pericratonic terrane (Gabrielse and Yorath, 1992). This area, the Cassiar Platform, is thought to have been part of the original Laurentian continent during the Paleozoic. However, it lies immediately to the west of the Rocky Mountain Trench–Tintina Fault, one of the longest and most signiﬁcant of the Cenozoic strike-slip faults that affected the Cordillera during the Cenozoic. Displacement along this fault is unknown, and it is therefore not known what the original relationship of the Cassiar Platform is to ancestral North America. Cecile and Norford (1993, p. 135) cited estimates of dextral displacement ranging from 400 to 750 km, the greater of which would place Cassiar Platform outboard from southern Alberta. Between the Tintina Fault and the craton is an area characterized by a complex array of Paleozoic basins and arches, including the large Selwyn Basin, in which deep-water (basinal) facies are widespread. Cecile et al. (1997)

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Figure 8 Below, restored cross-section through the craton and miogeocline of western Canada; above, two-dimensional numerical model of a £exural continental margin, based on the data and delithi¢cation procedures used for this particular margin. Note that (1) the gradual onlap of the craton with time; (2) the considerably greater thickness of strata in the continental-margin wedge than predicted by the £exural model; and (3) the thin unit of Middle to Upper Cambrian strata that extend for hundreds of kilometers into the cratonic interior. Inverted ‘‘v’s’’are well locations from which thickness data have been derived (Bond and Kominz, 1984).

suggested that this part of the continental margin may have been the lower-plate margin of a continental separation above a west-dipping detachment fault (Figure 7). The ‘‘pericratonic’’ character of the western part of the Paleozoic margin in this area may be attributed to the crustal attenuation and block-faulting characteristic of a lower-plate margin. As noted earlier, south of the 60th parallel the structural geology of the continental margin is consistent with that of an upper-plate margin, and Cecile et al. (1997) suggested that the Liard Line may have functioned as a transfer fault, across which the detachment fault reversed dip direction. A structural embayment in the northeast corner of the Selwyn Basin shows evidence of at least two episodes of rifting, one of Lower to Middle Cambrian Age and the other of late Early Ordovician to Middle Ordovician Age, indicating repeated episodes of crustal extension, similar to that which characterized the Grand Banks off Newfoundland during the Jurassic to Early Cenozoic (Cecile et al., 1997). North of the Cassiar Platform is the north–south-oriented Richardson Trough, a very long-lived crustal depression that functioned continuously as a deep-water environment for most of the Early and Mid-Paleozoic, from at least the Middle Cambrian to the Middle Devonian. This trough is located close to the pole of rotation around which the Canada Basin opened in the Cretaceous (Figure 21F, G in Chapter 1). Prior to that event, Richardson Trough was aligned with the Hazen Trough of the Canadian Arctic Islands, a similarly long-lived deep-water basin that was succeeded in the Late Paleozoic by the Sverdrup Basin (Chapter 13). It seems likely that Richardson Trough–Hazen Trough represents a very ancient lineament in the Laurentian crust that was established by rifting in the Late Precambrian or Early Paleozoic. A western extension of the Laurentian continent is the Yukon Platform, which is located to the west of Richardson Trough, and extends a short distance across the Alaska border, where it is truncated by faults associated with Cordilleran terrane accretion.

4. The Sauk Sequence and the Cambrian–Ordovician Shelf-to-Basin Transition Regional isopach maps (Cook and Bally, 1975) reveal a relatively continuous blanket of Middle and Upper Cambrian strata (Sauk II–III sequence) extending along the miogeoclinal margin, from Sonora, Mexico, through Nevada and British Columbia to Yukon. As discussed later, these strata onlap extensively onto the miogeocline and craton. They are truncated eastward by a Sub-Early Devonian (Sub-Kaskaskia) unconformity.

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4.1. The Kicking Horse Rim and Burgess Shale of the southern Canadian Rocky Mountains A transect across the Main Ranges of the Rocky Mountains on either side of the Trans-Canada Highway, just west of the Alberta–British Columbia border, provides the best exposed and most well-documented cross-section through the ancient Paleozoic continental margin. Geological reconnaissance through this region was facilitated by the construction of the Canadian Paciﬁc Railway, which was completed in 1885. In 1909, Charles Walcott, carrying out geological reconnaissance on horseback in the mountains above the railway, made the ﬁrst discovery of the exotic, proliﬁc, fauna in what came to be known as the Burgess Shale, of Middle Cambrian Age. His repeated visits to this area resulted in a very large collection being assembled at his home base, the Smithsonian Institution in Washington, DC, and the beginning of the formal description of the Paleozoic sedimentary rocks of the area (Walcott, 1928). Detailed re-examination of this area during the systematic mapping of the Rocky Mountains by the Geological Survey of Canada situated these fossils and their host strata in a modern stratigraphic and paleogeographic context (Fritz, 1971; Aitken, 1971). Researchers from the Royal Ontario Museum (Toronto), and elsewhere, have continued detailed studies of the Burgess Shale fauna. Aitken (1971) synthesized the stratigraphy of the Main Ranges of the Rocky Mountains just west of the Alberta–British Columbia border, an area lying mostly within Yoho National Park. He demonstrated that across a belt about 15 km wide, oriented northwest–southeast, Cambrian and Ordovician strata undergo a major facies change from cratonic, shelf facies, consisting predominantly of carbonates and mature clastics in the east, to a mainly deep-water clastic succession to the west (Figure 9). The belt within which the facies change takes place appears to have remained more or less stable in position from Middle Cambrian to Late Ordovician time. As noted earlier, Aitken (1971), termed this belt of facies change the Kicking Horse Rim (Figure 1). The entire Cambrian–Ordovician section totals about 3,400 m in this area (Aitken, 1993a). Further, the stratigraphy of the Cambrian section suggests the former existence of three major paleogeographic belts: 1) An inner detrital facies, deposited toward the center of the craton, and characterized by shallow-water sandstone, siltstone and shale. Sandstone increases in importance toward the shoreline on the inner margin of the belt; glauconite is a common accessory. 2) A middle, carbonate–shoal facies. Clastic units are rare, but the carbonates may contain grains of clay and quartz sand and silt. 3) An outer detrital belt of thin-bedded sandstone, mudstone and carbonate. The existence of these three broad facies belts was ﬁrst suggested from study of the Cambrian section in the Great Basin of Nevada (Robison, 1960; Palmer, 1960), and found wide application in the study of the Lower Paleozoic rocks of the Canadian Rocky Mountains. Figure 10 illustrates one version of this model as applied to the Middle Cambrian strata of the Rocky Mountains. The inner detrital belt, subdivided into an inshore basin and a shoreline area, corresponds to the interior of the craton, beyond the scope of this chapter. The Kicking Horse Rim corresponds to the carbonate–shoal area, and the open basin of Figure 10 corresponds to the outer detrital belt, an area that had also been termed the Robson Trough in earlier studies. Although this latter term no longer seems appropriate, with the recognition that the deep-water area is an ancient continental margin, open to the Panthalassa Ocean to the west, the terms Robson Basin and Columbia Basin have been used for the Early Paleozoic areas of deep-water sedimentation in British Columbia that are immediately adjacent to the Laurentian Margin (Figure 1). Aitken (1971) demonstrated that in places the facies change from the carbonate to the outer detrital belt is a gradual one, as if down a ramp; in places it is marked by major slumps and slides; and in the vicinity of the famous Burgess Shale fossil localities, the facies change takes place across a distinct fossil escarpment about 200 m high (Figure 11). The intense interest in the Burgess Shale fauna has led to detailed mapping and repeated reexaminations of this area by numerous geologists (e.g., McIlreath, 1977; Aitken and McIlreath, 1982; Fletcher and Collins, 1998). This work has conﬁrmed that the escarpment was a contemporary feature of the submarine landscape during the Middle Cambrian. Careful biostratigraphic study (primarily of the trilobites), commencing with Fritz (1971), demonstrated that the rocks beneath the Burgess Shale, within the deep basin, the Takakkaw Limestone Tongue, are the same age as the rocks of the Cathedral Rim, at the top of the Cathedral Limestone Formation. The Burgess Shale therefore consists of somewhat younger Middle Cambrian sediment, derived by transport of ﬁne-grained clastic detritus across the edge of the escarpment and then banked up against, and eventually burying, the scarp face. There is very little evidence of direct derivation of sediment from the scarp itself. The Wash Limestone Member contains thin carbonate debris-ﬂow beds that probably represent a local collapse of the scarp face. The rocks at this margin consist of reef-ﬂat fenestral, stromatolitic and thrombolitic carbonates with oolites and grainstones, but they are pervasively dolomitized and have lost much of their primary

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Figure 9 Changes in thickness and facies of the Cambrian--Ordovician section of Yoho National Park, Canada, as summarized by Aitken (1971).

fabric. At the time the Burgess Shale accumulated, the Cathedral Rim was probably not an actively growing reef construction (Aitken, 1989). The most popular model for the deposition of the Burgess Shale is that it represents a very quiet-water environment in the lee of the scarp face. This would account for the preservation of the unusual and abundant soft-bodied fauna. There is no evidence for deposition by turbidity currents, in the form of graded bedding or sole markings. However, there is also virtually no evidence of bioturbation of the soft sediments, even within the beds where the proliﬁc fauna has been studied in detail. This is one of the lines of evidence that led Gostlin and Miall (2005) to dispute the conventional interpretation, that the organisms lived in the deep basin, where their remains are now found. Instead, it was suggested that the organisms lived in the shallow waters above and behind

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Figure 10 Paleogeographic model for the western craton margin during Middle Cambrian time, in the area of the Rocky Mountains of Yoho Park, Canada, showing the major facies belts. Grand Cycles are discussed in the text (Aitken, 1989). Modi¢ed from Aitken (1978).

the Cathedral Rim, and were swept over the scarp edge by occasional storms. They are therefore, for the most part, not preserved in their original living positions. Beyond (west of) the classic Burgess Shale locations, in the Columbia Basin, outcrops of the basinal facies, The Chancellor Group, are sparse. Aitken (1993a, p. 111) reported that the group ‘‘consists of shale, laminated siltstone, ribbon-bedded, more or less argillaceous and silty lime mudstone, and limestone-shale couplets. Debrisﬂow breccias and large olistoliths are present at several levels.’’ The Group totals about 1,600 m in thickness.

4.2. Northern British Columbia To the northwest of the classic Yoho Park locations discussed earlier, in the Robson Basin of northern British Columbia, a conformable succession of Upper Precambrian to Lower Cambrian quartzites is present. These have been assigned to the Hamill Group in southeastern British Columbia, and to the Gog Group, in areas along strike to the northwest. The latter is up to 2,200 m thick. The Middle and Upper Cambrian section in this area is similar to that comprising the middle carbonate–shoal facies of areas further south. Mixed shallow-water carbonate and ﬁne clastic units predominate. The total Cambrian–Early Ordovician section in the Robson Basin reaches 6,300 m in thickness (Fritz et al., 1992). A description of the passage from miogeoclinal to deep-water facies is not available for this area.

4.3. Yukon and northwest territories Strata of earliest Early Cambrian Age are not present in Richardson Trough and Yukon Platform (Fritz et al., 1992). Sedimentation in Richardson Trough commenced with the Illtyd Formation, a pure carbonate more than 600 m thick, succeeded by the Slats Creek Formation, a marine to nonmarine clastic unit, up to 1,500 m thick, containing volcanics. Extensional faulting appears to have established the identity of the trough at this time. The Slats Creek is followed by the Road River Formation, a unit up to 3 km thick that appears to represent virtually continuous, deep-water sedimentation in local basins of a predominantly muddy facies for tens of millions of years, until the Mid-Devonian. Carbonate deposition took place on adjacent platform areas. On Cassiar Platform, carbonate sediments with archeocyathid bioherms of Lower Cambrian Age are present. The Middle Cambrian is absent and strata of Late Cambrian Age are poorly known.

4.4. Great Basin: Nevada, Utah, Idaho A terrigenous detrital sequence spanning the Precambrian–Cambrian boundary is widespread in the Great Basin (Stewart and Suczek, 1977; Poole et al., 1992). A quartzite unit occurs at the base, consisting of cross-bedded sandstone, commonly conglomeratic, interstratiﬁed with siltstone and argillite. Units of limestone and dolostone are also present, locally reaching 600 m in thickness. The clastic deposits contain algal remains and (in southern

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Figure 11 The relationship between the Burgess Shale and the Cathedral Limestone in the vicinity of the original ‘‘Walcott Quarry,’’on the £ank of Mount Field, immediately to the north of the Trans-Canada Highway, near Field, Yoho National Park, British Columbia (Fletcher and Collins, 1998).

Nevada and eastern California) archeocyathids, and are characterized by shallow-water sedimentary structures, including bimodally oriented crossbedding, ﬂaser and lenticular bedding, and are interpreted as tidal in origin. West of the continental margin, in Nevada, the sequence is locally more than 6 km in thickness. It thins to the east, and a 300 m isopach was used by Stewart and Suczek (1977) to differentiate a ‘‘cratonal’’ facies. Middle and Upper Cambrian strata constitute a ‘‘carbonate sequence’’ in the Great Basin. This is a westerly thickening succession, ranging from less than 200 m within the miogeocline, to as much as 1,500 m in western

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Nevada. The facies belts erected by Palmer (1960) and Robison (1960) for this area can be traced south from Canada into the western United States and northern Mexico, although the inner detrital belt is only present well within the miogeocline, in Montana and Arizona (Stewart and Suczek, 1977). The middle carbonate belt consists of a variety of shallow-water limestones, containing stromatolitic boundstones, oolites, oncolitic and pellet grainstones and wackestones. The outer detrital belt shows a westerly increase in shales, graded bedded grainstones and wackestones, slump structures and intraformational conglomerates, and the local presence of chert. This indicates a gradation westward into basinal slope facies.

5. Middle Ordovician–Early Devonian (Tippecanoe Sequence) Strata of Middle Ordovician to Early Devonian Age (Tippecanoe Sequence) are rather more patchily distributed than those of the underlying Sauk Sequence, owing to epeirogenic warping (Figure 12). The West Alberta Ridge or Arch was active during the Mid-Paleozoic, from the Middle Ordovician to the Early Devonian. Cecile and Norford (1993) suggested that the Peace-Athabasca Arch (shown with its alternative name, Peace River Arch in Figure 12) was also active as a low-relief sediment source. Rocks of this age are absent over most of western Alberta and adjacent areas of British Columbia. The two arches effectively separated the continental margin into two broad depositional areas, (1) northern Canada, from northern British Columbia to Yukon and (2) the Great Basin and a miogeoclinal embayment to the Williston Basin of the cratonic interior. Aitken (1993b) attributed the West Alberta Arch and other mildly active positive features to changing patterns of intraplate stress acting across the craton in response to plate-tectonic forces acting at the continental margins. Cecile et al. (1997) interpreted this arch as an uplift associated with an upper-plate continental margin above an east-dipping detachment fault that developed during Late Precambrian to Early Paleozoic continental separation. Sedimentation was presumably continuous along the deeper parts of the continental margin, between Idaho and British Columbia, but these rocks have not been preserved (or have yet to be identiﬁed) within this area of the Cordilleran Orogen.

5.1. Northern Canada The Mackenzie Mountains area of the Northwest Territories, and Yukon Territory, consisted of a large number of platforms, basins and embayments which underwent repeated epeirogenic movement during this period. Stratigraphic thicknesses are very variable; they exceed 900 m in the Liard Depression, which is described as a basin that underwent abnormally rapid subsidence during this period (Cecile and Norford, 1993, p. 143). The Liard Depression, Richardson Trough and the Selwyn Basin are characterized by shale, cherty shale, and dolomitic mudstone and siltstone. The eastern margin of this facies (Figure 12) deﬁnes an approximate boundary to the deeper-water, basinal margin of the Laurentian continent. The broad extent of Tippecanoe deposits in Yukon and Northwest Territories, north of the Liard Line, is noteworthy. They extend eastward across the entire craton to the margin of the Precambrian Shield. This is part of the evidence suggested by Cecile et al. (1997) for their interpretation of this part of the continental margin as a lower-plate margin above a west-dipping detachment fault. Typical of the many platform carbonate units that have been mapped in this area is the Mount Kindle Formation of the Mackenzie Mountains. This unit, consisting of thick-bedded, fossiliferous dolostone, ranges between 100 and 400 m in thickness. Many other units, too numerous to describe in this brief chapter, have been erected for the localized and variable facies belts of the northwest. Cecile and Norford (1993) have described the Lower Paleozoic platform succession as consisting of a suite of three transgressive–regressive cycles. Cycle B, of which the Mount Kindle Formation constitutes the greater part, corresponds approximately to the Tippecanoe Sequence, although there appears to be no evidence of the endOrdovician regional unconformity within this unit that elsewhere separates the Tippecanoe Sequence into two subsequences (Figure 2).

5.2. Great Basin Cyclic sedimentation patterns and the presence of local widespread disconformities indicate a similar pattern of regional tectonic and global sea-level control in Nevada to that of Northern Canada. For example, Figure 13 shows a cross-section through the Middle Ordovician section from Utah to Nevada. ‘‘The Whiterockian carbonate shelf of western Utah, Nevada and southern California developed during a single ofﬂap–onlap cycle, 12 m.y. in duration, while most of the North American continent was exposed to subaerial weathering’’

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Figure 12 Isopach of the Tippecanoe Sequence (Middle Ordovician to Early Devonian), (adapted from Cook and Bally, 1975; Sloss, 1988; Aitken, 1993a). The extent of shale facies is from Miller et al. (1992).

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Figure 13 Middle Ordovician facies relations from the miogeocline of Utah to the continental margin of southwestern Nevada and adjacent areas of California (Poole et al., 1992).

(Poole et al., 1992, p. 22). Graptolitic back shales of the Kanosh Shale were deposited in Utah, while the shallowbank and tidal-ﬂat carbonate deposits of the Antelope Valley Limestone were forming to the west. Banks of sponges and algae appear to have formed barriers, behind which the graptolitic muds were deposited. The Eureka Quartzite, which follows, consists of great thicknesses of relatively pure quartz sand thought to have been derived by erosion of pre-existing quartz sandstone deposits within the cratonic interior. Upper Ordovician strata of the continental-slope and rise facies are preserved within the Roberts Mountains Allochthon, and are thought to have been thrust eastward from an originally off-shelf setting during the Antler Orogeny. They consist of strongly deformed shaly, siliceous rocks containing planktonic and nektonic fossils. A widespread dolostone facies up to 300 m thick (Red River, Bighorn formations) extends across the miogeocline to the east. Silurian and Lower Devonian strata are well represented in the Great Basin (Poole et al., 1977), thinning eastward onto the ﬂanks of the Transcontinental Arch. On the miogeocline margin, Silurian rocks total as much as 600 m in thickness, and consist mainly of shallow-water dolomites and sandy dolomites. A distinct facies change into slope deposits has been mapped near the east edge of the Roberts Mountains Allochthon. There, the shelf dolomites become cherty westward, and grade laterally into the laminated limestone and silty limestone of the Roberts Mountains Formation. Sparse fossils of a shelly fauna are present in the shelf dolomites. In the slope deposits fossils are more abundant, and consist of both the shelly and graptolitic facies. Lower Devonian strata also contain a distinct thickness and facies change passing from the craton margin to the continental slope. Cratonic sedimentary rocks are predominantly nearshore and shallow subtidal to intertidal dolomites and sandy dolomites. In the southern part of the shelf a cherty, argillaceous dolomite occurs near the top of the Lower Devonian. Westward, these rocks grade into slope deposits of thick-bedded detrital limestone and thin-bedded sandy to argillaceous limestone with subordinate beds of mudstone and dolomite. The section thickens in the same direction, from 100–300 m on the shelf, to locally as much as 1,000 m on the outer shelf and slope.

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6. Grand Cycles The Cambrian–Ordovician stratigraphy of Laurentia is strongly cyclic. The concept of the Grand Cycle was ﬁrst proposed based on studies in Nevada (Robison, 1960; Palmer, 1960), but the ideas were more fully explored and developed in application to the Canadian portion of the continental margin, along the Kicking Horse Rim (Aitken, 1966, 1978), and have now been widely applied to cyclic Cambrian deposits throughout Laurentia (e.g., Cowan and James, 1993; see also Chapter 3). Eight Grand Cycles have been identiﬁed in the Cambrian to Middle Ordovician stratigraphy of the Canadian Rocky Mountains (Aitken, 1993a, p. 109). Five of these are shown in Figure 14. Cycle 1, of Early Cambrian Age, corresponds approximately to the Sauk-I Sequence of Sloss (1988). The next four cycles, of Middle Cambrian and earliest Late Cambrian Age, represent a higher-order cyclicity within the Sauk-II Sequence. The top three cycles (Upper Cambrian–Middle Ordovician) overlap from the Sauk III into the basal Tippecanoe I Sequence. The boundaries of these two younger Sloss subsequences do not appear to correspond to the Grand Cycle boundaries, suggesting that eustatic sea-level change does not provide a complete explanation of the cycle driving processes, at least for the upper three cycles. Aitken (1978) identiﬁed two broad types of Grand Cycle: ‘‘Stephen-type cycles’’ were deposited when the inshore basin behind the Kicking Horse Rim was conﬁned behind a narrow and discontinuous rim of intertidal to supratidal carbonate–shoals, permitting easy tidal exchange between the inshore basin and the open ocean, and ‘‘low-energy’’ conditions in the inshore basin. ‘‘Sullivan-type cycles’’ were deposited when the inshore basin was conﬁned behind a wide, practically unbreached carbonate–shoal complex. Limited tidal exchange across the shoal complex combined with favorable orientation and scale of the inshore basin resulted in locally high tidal range and consequent high tidal energy at the continental margin. The classic ‘‘Stephen-type’’ cycle commences with ‘‘ﬂaggy’’ lime-mudstones, with mudstone partings of the basal Stephen Formation. Feeding burrows and trilobite fossils are abundant. There are minor internal shale– limestone cycles a few meters thick. Facies trends suggest a gradual shallowing of the depositional setting with time. In the middle of the cycle, the proportion of shale increases, and there are minor rippled laminae of quartz siltstone and calcisiltite. Brachiopod fragments are common. Interbedded units of thrombolites and stromatolites appear, together with lime-mud partings with ripples. The Stephen–Eldon contact is transitional, with shales becoming rarer, while limestones become predominant. Higher in the Eldon Formation a mottled facies with dolomitized borrows becomes characteristic. Occasional interbeds of pellet and intraclast grainstone occur. This type of cycle is interpreted as the product of marine transgression. Initially a peritidal carbonate–shoal complex existed at the Kicking Horse Rim, but this was gradually drowned with increasing turbulence, permitting the transportation of detrital material westward into the deeper parts of the basin. Sea levels then gradually fell, again, permitting a re-establishment of the carbonate rim. The type example of the Sullivan-type cycle begins with an abrupt transition from the grainstones of the Waterfowl Formation, into the shale of the Sullivan Formation. This lithology is somewhat calcareous and sparsely fossiliferous, with fragments of trilobites, brachiopods and echinoderms. Interbedded with the shale are

Figure 14 Schematic stratigraphic cross-section of the Cambrian rocks of the western continental margin, extending from Jasper National Park southeastward to Ban¡ National Park, and then westward across the Kicking Horse Rim (Fritz et al., 1992). Six of the Grand Cycles of Aitken (1966, 1978) are highlighted, but the top of Cycle 5, which is of Upper Cambrian Age (Lyell Formation), is not shown. Reproduced with the permission of the Minister of Public Works and Government Services Canada, 2007 and Courtesy of Geological Survey of Canada.

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units of dolomitized skeletal grainstone and packstone, and cryptalgal limestone. There is an upward transition into the predominantly carbonate Lyell Formation, most of which consists of shoaling-upward carbonate subcycles constituting packages of carbonate conglomerate, cryptalgal laminate and calcisiltite laminate. Aitken (1978) suggested that the Lyell Formation represents a carbonate–shoal complex some 400 km across, with oolitic carbonate sands accumulating in tidal banks comparable to those developing at the present day in the Persian Gulf and the Bahamas. Both cycle types are interpreted as the product of cycles of eustatic sea-level rise and transgression, resulting in the migration or drowning of the carbonate–shoal at the margin of the craton. However, as noted earlier, detailed biostratigraphic correlation does not support a precise correspondence between the younger three cycles and the Sloss sequences, suggesting that in the southern Rocky Mountains accommodation changes may have also been inﬂuenced by other processes, such as epeirogenesis driven by dynamic topography.

7. Lower to Upper Devonian (Kaskaskia-I Sequence) By Early Devonian time, a clear distinction can be made in the Great Basin, between continental margin strata, deposited on the shelf or slope of Laurentia, and arc-related rocks of the approaching Antler Arc (Poole et al., 1977). Figure 15, which summarizes the thickness distribution and facies of Mid-Devonian strata, is provided here as an indication of the general paleogeographic conﬁgurations during Kaskaskia-I sedimentation. On the western continental margin, the initial Kaskaskia transgression during the Early Devonian, generated a cratonic seaway that extended all along the western margin of Laurentia (Cook et al., 1975). The West Alberta Arch was transgressed by Middle Devonian time, and ceased to be recognizable as a distinct cratonic arch. Meanwhile, however, the Peace River Arch became active. Some evidence for its presence as a modest sediment source, exists in Late Precambrian and Early Paleozoic strata, and by the Mid-Devonian, a distinctive basal clastic facies was being deposited in erosional hollows across the uplifted Arch (O’Connell, 1994). The Ancestral Uinta Uplift was also active as a clastic sediment source during the early Late Devonian (Poole et al., 1977). This uplift and the Peace River Arch are transverse elements; that is, they are oriented at a high angle to the continental margin. More or less simultaneous movement on these two unusual tectonic elements may indicate a response to a temporary change in intracontinental intraplate stress patterns. Slope-and-basin deposits are well represented in Yukon, where the Selwyn Basin encompassed a large area of thinned, subsided and downfaulted continental crust. Across the Tintina fault lies the Cassiar platform, a pericratonic terrane of originally Laurentian afﬁnities but probably displaced northwards hundreds of kilometers from its original location. Along the margin between the Peace River Arch and the northern Great Basin little evidence of slope-and-basin facies is preserved, this facies belt having been metamorphosed and upthrust, or eroded during Cordilleran tectonism (this is discussed later).

7.1. Northern Canada Kaskaskia strata have been subdivided by Fritz et al. (1992) and Gordey et al. (1992) into sequences, corresponding to ‘‘allogroups,’’ as deﬁned by the North American Commission on Stratigraphic Nomenclature (1983). Although this provides for convenience in classiﬁcation and description, the actual complexity of the stratigraphy in this large area suggests that the region was affected by many local episodes of tectonism which have obscured large-scale trends. A summary of the stratigraphy and broad facies relationships is shown in Figure 16. Miogeoclinal sedimentation on the Mackenzie Platform was characterized by a succession of largely carbonate units, consisting variously of dolomite, limestone, sandy limestone, siltstone and shale. Reef barriers formed at several different times during the Devonian (the extensive, petroleum-producing Devonian reefs of the Alberta Basin are beyond the scope of this chapter). Evaporites formed behind a shelf-margin barrier at several times (e.g., Bear Rock Formation, Elk Point Group). Dramatic facies changes into deeper-water, predominantly clastic facies, have been mapped along the eastern margin of Selwyn Basin. Morrow and Geldsetzer (in Fritz et al., 1992, p. 204) noted that Lower Devonian (Siegenian to Emsian) strata consist of black, graptolitic shale, crinoidal limestone, argillaceous limestone, dolomitic sandstone and units of ﬁne-grained quartzite up to 100 m thick. Carbonate and siliciclastic units were deposited as debris ﬂows and turbidity-current deposits derived from the platform to the east. In Selwyn Basin and Richardson Trough, calcareous shale, limestone and dark, siliceous shale of the Road River Formation and other, similar units, were deposited through much of the Early and Middle Devonian.

Figure 15 Schematic paleogeography and generalized facies trends, Middle Devonian of the western Laurentian Margin, corresponding to the middle part of the Kaskaskia-I Sequence. Position of the continental margin is generalized --- in northern Canada areas of uplift and subsidence and of deep- and shallow-water sedimentation underwent regular change owing to active local tectonism. Isopachs are shown where data are consistent enough to indicate local trends. Broad facies characteristics are indicated by black and white ornamentation. Areas lacking ornamentation are areas characterized mainly by shallow-water carbonate sedimentation. Data for the United States is from a map detailing the Frasnian and Lower Fammenian (lower Upper Devonian) paleogeography of the Great Basin, from Poole et al. (1992, Plate 3). Data for Canada are from maps showing thickness and facies trends for the Mid-Givetian to Famennian, from Gordey et al. (1992). Locations of the two cross-sections of Figure 16 are shown.

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Figure 16 Stratigraphic cross-sections through the Devonian system of western Canada. Gray shading indicates predominantly clastic sediments. Locations are shown in Figure 15 (from Morrow and Geldsetzer, in Fritz et al., 1992). Reproduced with the permission of the Minister of PublicWorks and Government Services Canada, 2007 and Courtesy of Geological Survey of Canada.

Between the late Middle and the early Late Devonian, carbonate sedimentation largely ceased in northern Canada. Tectonism in northwest Yukon and adjacent areas of Alaska, including granitic intrusion and uplift, generated uplifted orogens, from which sediment was shed south and east across the entire Selwyn Basin and Mackenzie Platform area. The northern craton was also uplifted and underwent exposure and erosion at this time. The dominance of clastic sedimentation began in Yukon in the early Middle Devonian. In the Mackenzie Platform area, carbonates of the Hume Formation were succeeded in the Mid-Devonian by the calcareous shalelimestone succession of the Hare Indian formation, then by black organic shale of the Canol Formation, and ﬁnally by the thick, ﬁne-grained clastic succession of the Imperial Formation. This unit, which reaches thicknesses of more than 2 km at the northern margin of the Mackenzie Platform, consists of shale, siltstone, ﬁnegrained sandstone and minor limestone, derived from cratonic sources to the east (and possibly ultimately from Caledonian sources on the eastern margin of Laurentia — see Chapter 17). It was deposited in shallow-marine nearshore to offshore environments. To the west and north, in the area of Richardson Trough, the Imperial formation consists of deeper-water deposits, predominantly turbidite sandstone, derived from orogenic sources in NW Yukon and Alaska. By the early Late Devonian, clastic sedimentation had spread to southern Alberta, where it constitutes the Ireton Shale, an important organic source rock and reservoir cap rock (Gordey, in Gordey et al., 1992).

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A succession of Upper Devonian to Mid-Mississippian rocks more than 700 m thick is present on the Cassiar Platform (Gordey, in Gordey et al., 1992). These consist mainly of siliceous mudstone and ﬁne-grained siltstone. Debris-ﬂow conglomerates are also present, containing abundant clasts of siliciﬁed and sandy carbonates thought to have been derived from the nearby miogeocline. The facies suggest a rapid subsidence of the platform, and the indication of a nearby miogeoclinal clastic source suggests that the platform at this time was functioning as a rift basin adjacent to the continental margin.

7.2. Peace River Arch This craton–margin tectonic element has been characterized by anomalous tectonic episodes relative to the adjacent craton for much of the Phanerozoic (O’Connell, 1994). It became paleogeographically signiﬁcant during the Middle and Late Devonian, when it was uplifted and subject to signiﬁcant erosion. A clastic unit, known as the Granite Wash blanketed the uplift, and interﬁngered with marine carbonate, shale and evaporite units to the north, east and south. The Granite Wash is up to 100 m thick over the crest of the arch, and is composed of material derived from erosion of the uplift, predominantly Precambrian granitic and metasedimentary rocks, indicating deep erosion of the earlier Paleozoic cover. The deposit ﬁlls fault-bounded rift basins across the arch. At the margins of the arch it interﬁngers with estuarine and ﬂuvial sand bodies and the arch is encircled by a series of fringing and patch reefs spanning the Middle to Early Upper Devonian. Together with the Tathlina Uplift and the West Alberta Arch these areas of craton–margin uplift served as barriers to marine circulation across the continental interior during the Middle Devonian, leading to restrictedmarine environments and the deposition of the Elk Point and other widespread evaporite deposits across eastern Alberta and Saskatchewan. Early in the Late Devonian the Peace River Arch became passive, and was onlapped by the marine rocks of the miogeocline.

7.3. Ancestral Uinta uplift This is another unusual east–west oriented tectonic element, in that it shed synorogenic sandstone and conglomerate during the early part of the Late Devonian from a local uplift on the inner shelf into what was otherwise a broad miogeoclinal sea characterized by deposition of carbonates, now mostly dolostones (Poole et al., 1992). Transport directions recorded in the Stansbury Formation, are predominantly eastward.

7.4. Great Basin As noted earlier, most of the miogeocline margin of the Great Basin area was undergoing shelf carbonate sedimentation during the Mid- to Late Devonian (Figure 15), although the distribution of rocks of this age is now limited owing to deep pre-Mississippian erosion (Poole et al., 1992). The Pilot Shale (Late Upper Devonian), of eastern Nevada, is interpreted as a ‘‘protoﬂysch’’ reﬂecting the early inﬂuence of the approaching Antler Orogen (Poole et al., 1977). It consists of siltstone, carbonaceous siltstone and mudstone with interbeds of turbidite and debris-ﬂow deposits. Sedimentation of passive-margin type on the ancient continental margin in the Great Basin area was brought to a close by the Antler Orogeny in latest Devonian or earliest Mississippian time. This and other arc-related tectonic episodes and related sedimentary units are discussed in Chapter 11.

8. Mississippian Arc Collisions and Termination of Parts of the ‘‘Passive’’ Laurentian Margin (Kaskaskia-II Sequence) The Antler Orogeny emplaced the Roberts Mountains Allochthon on the continental margin, above a thrust belt extending from southeastern California to central Idaho, in the latest Devonian or Early Mississippian, permanently changing the tectonic character of the western continental margin of that area (Chapter 11). Within Canada, the change to a convergent margin took place much later. The evidence suggests that at least part of the margin, that part lying within eastern British Columbia, remained a ‘‘passive’’ margin until the Triassic or Early Jurassic. However, further to the north, the Kootenay Terrane of west–central Yukon and adjacent areas of Alaska, has been interpreted as a west-facing arc of Upper Devonian and Early Carboniferous Age (Richards, 1989). The name Prophet Trough has been assigned to the belt of downfaulted continental margin rocks of latest Devonian

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and Carboniferous Age extending from southern Yukon to northern Idaho, and possibly linking up with the Antler foreland basin (Richards, 1989; Gordey et al., 1992). Rift faulting in these rocks could be extensional faults related to passive-margin or backarc rifting, or it could indicate crustal downwarping as a result of contractional loading by the approaching Kootenay Arc. Within this model, the pericratonic Cassiar Platform would be interpreted as a retroarc foreland basin or a backarc basin, depending on whether the arc was contractional or extensional. Gabrielse and Yorath (1992, p. 691) noted that no contractional structures comparable to those associated with the Roberts Mountains Allochthon have been identiﬁed in northern Canada, but that ‘‘it is conceivable that the rifting, volcanism, uplift and sedimentation in the northern Cordillera was linked to tectonism to the west in the Kootenay terrane.’’ Rocks of Mississippian to Early Pennsylvanian are well represented along the western margin. Figure 17 provides an isopach map for the Lower Mississippian. Contours cannot be drawn for some areas of occurrence, such as the Prophet Trough, because of structural complexities.

8.1. Great Basin A foredeep developed in front of the Antler Orogen on the outer continental shelf. Figure 17 shows the maximum areal extent of this foreland basin, which was reached during the Late Mississippian (Miller et al., 1992, Plate 4C). Shallow-water carbonates interﬁngered with siliciclastic deposits derived from the emerging Roberts Mountains Allochthon. Deep-water slope and submarine-fan deposits prograded cratonward (eastward) during the Early Mississippian, and in the Late Mississippian graded up into deltaic deposits as the foredeep ﬁlled with sediment and became shallower. By this time, it appears that contractional tectonism of the Antler Orogen had ceased (Miller et al., 1992). On the craton, shallow-water limestone and sandstone accumulated, with a depocenter in the area of the present Uinta uplift. A regression commenced in the Mid-Mississippian, terminating the Kaskaskia Sequence, and much of the North American Craton became exposed as a vast karst plain.

8.2. Western Canada Carbonates and shales of Late Devonian and Mississippian Age, including such units as the Palliser and Rundle formations, constitute some of the most spectacular mountain ranges of the Rocky Mountains in Banff and Jasper national parks (Price and Mountjoy, 1970). These are almost entirely cratonic in origin. In the western main ranges, close to the Alberta–British Columbia border, the basal Banff Formation, of Early Mississippian Age, grades westward and stratigraphically downward into the Besa River Formation, of latest Devonian and earliest Mississippian Age (Gordey et al., 1992). The Banff Formation is a limestone, dolomite and silty carbonate unit of shelf-margin and slope origin, and the Besa River Formation consists of basinal shale and dolomitic shale. A similar transition has been identiﬁed at several places along the eastern margin of the Prophet Trough, for example near the 60th parallel and in the Eagle Plain. These shelf-slope-basin assemblages are almost everywhere progradational, with shallower water facies gradually advancing basinward (westward). In the Mississippian, the Peace River Arch underwent collapse, and became a basin, termed the Peace River Embayment (Figure 17). From here northwards the Rundle Group, of late Early to Mid-Mississippian Age, shows a westward transition from a shelf and slope carbonate facies into a 200 m-thick unit of spicule-rich carbonate grainstones, called the Prophet Formation, and thence into the basal Besa River shales. Large-scale submarine erosional channels are common in these slope deposits. This is overlain by the thick deltaic sandstone–shale succession of the Stoddard Group and Mattson Formation, which ﬁlled the Peace River Embayment and the Liard Basin. The Cassiar terrane contains a succession about 400 m thick of Mississippian deposits, consisting of a lower sandstone–shale succession and an upper argillaceous, thick-bedded limestone. As noted earlier, there is no clear evidence for the approach and collision of the Kootenay Arc with Laurentia, although this has been postulated by several authors (e.g., Ricketts, 1989). The classic characteristics of a foreland basin — a detrital sedimentary wedge thickening away from the craton and showing evidence of derivation from the west, have not been claimed for Prophet Trough (in contrast to the very clear evidence for this tectonic setting for the Antler foreland basin of Nevada, see Chapter 11). However, an alternative model for this basin as a backarc basin behind an extensional or neutral arc (in the terminology of Dewey, 1980) would appear to be consistent with the evidence. The Kaskaskia Sequence was terminated by a widespread regression and deep erosion over much of the craton (Henderson, 1989).

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Figure 17 Distribution of strata of ‘‘pre-Chesterian’’age on the western continental margin, showing major contemporaneous paleogeographic elements. The time interval used in this map corresponds to the lower half of the Mississippian (Tournaisian and lower Visean). Isopachs: Cook et al. (1975), Additional Canadian data: Gabrielse and Yorath (1992), US data, Miller et al. (1992).

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9. Pennsylvanian–Permian (Absaroka-I and II Sequences) ‘‘Antler overlap basins’’ developed during the Pennsylvanian and Permian of Nevada (Miller et al., 1992). Little evidence of an ancient continental margin is preserved along the western edge of the craton, from northern Nevada through Idaho. The Oquirrh Basin of northwest Utah contains more than 6 km of deep- to shallowwater carbonates and clastic of Pennsylvanian and Permian Age (Rich, 1977; Burchﬁel et al., 1992), but the basin appears to be the result of craton–margin subsidence, rather than an embayment in the continental margin. Development of the Oquirrh Basin may be related to intracratonic tectonism caused by the Late Paleozoic collision of Gondwana and Laurentia (see Chapters 7, 8). In Canada, strata of this age are best represented in thrust slices along the Rocky Mountains of British Columbia, and in the subsurface in the northeast part of this province and adjacent areas of west–central Alberta. Tectonic elements that were established during the Mississippian continued with little change through the Pennsylvanian and Permian. A basin extended from northern Yukon to northern British Columbia in much the same position as the earlier Prophet Trough, but was named the Ishbel Trough by Henderson (1989). Peace River Embayment persisted as an area of subsidence and clastic sedimentation through the Pennsylvanian. Sediments of Pennsylvanian and Permian Age are predominantly siliciclastic, in contrast to the thick and widespread carbonates of Mississippian and older strata. This is probably in part a reﬂection of the northward drift of Laurentia toward higher latitudes that were less favorable for carbonate biogenesis. Henderson (1989), noted that Permian brachiopod faunas have boreal afﬁnities, and are interpreted as temperate in origin. Absaroka-I and Absaroka-II have been divided into ﬁve sequences in Canada, separated by regional unconformities. These mostly consist of mixed clastic–carbonate successions on the shelf, but show westward facies transitions into slope facies composed of siltstone, dolomitic siltstone and shale (e.g., Johnston Canyon and Kindle formations, of Permian Age). However, as is the case with the underlying Kaskaskia Sequence, there is no clear evidence of a continental margin, in the sense of deep-water sediments resting on thinned continental crust facing an open ocean. Any remnants of such a margin have been destroyed by Cordilleran Orogeny, including thrust faulting and uplift and, during the Cenozoic, major right-lateral strike-slip faulting. A tectonic interpretation of the British Columbia continental margin is discussed later.

10. Triassic–Jurassic: Termination of the ‘‘Passive’’ Continental Margin Triassic rocks (Absaroka-III Sequence) are thick and well exposed in the thrust belts of the Rocky Mountains, from the US border northwest to the 60th parallel. They thicken westward from an erosional edge in west–central Alberta to a maximum of more than 1,200 m in the foothills belt of northeast British Columbia (Gibson, in Gordey et al., 1992). The isopach pattern, paralleling the ancient continental margin, suggests that the thickening is the result of subsidence of the continental margin, but there is no evidence of any distinctively deep-water facies, and as is the case with the underlying Devonian to Permian succession, it would appear that the true continental margin, that is, a deep-water basin developed over thinned continental crust, is now part of the deformed Omineca belt, to the west. In Banff National Park Triassic rocks comprise the Spray River Group, a mixed succession of ﬁne-grained clastics and sandy and silty carbonates. Some of the clastics are interpreted as thin-bedded turbidites, and would appear to constitute the last major suite of siliciclastic sedimentary rocks derived by erosion of the cratonic interior prior to the emergence of the Cordilleran Orogen and the reversal of transport directions into the newly formed Western Interior Seaway. Models of Cordilleran Orogeny for southern British Columbia suggest that the Middle Cambrian to Permian continental margin of this area had, by Late Paleozoic time, evolved into the continental ﬂank of a backarc basin (Figure 18). During the Jurassic, the west-facing Nicola Arc converged against the continental margin, resulting in obduction of continental ﬂakes, delamination of continental crust from its basement roots and the development of tectonic wedges that thrust eastward. The ancient miogeocline underwent regional metamorphism and deformation, generating what is now the Omineca Belt, comprising an imbricated succession of folded thrust sheets (Gabrielse and Yorath, 1992; Price and Monger, 2003). The ﬁrst phase of uplift of what became the Omineca Belt in the Late Jurassic generated the source for the ﬁrst westerly derived detritus to be shed from the newly created Cordilleran Orogen, and this marked the initiation of the Western Interior Seaway (Chapter 9). The evolution of the continental margin southward through Idaho and into northern Nevada may have followed a similar pattern to that in southern British Columbia, but the evidence is very sparse, owing to intense deformation during the Mesozoic (Burchﬁel et al., 1992). In northern British Columbia the continental margin

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Figure 18 Tectonic model for the evolution of the continental margin of southern British Columbia during the Jurassic. Location of cross-section is shown in Figure 1. Diagram A shows the margin during the Early Jurassic. The miogeocline, as shown in this panel, represents the continental-margin sedimentary rocks of Cambrian to Triassic Age that are described in this chapter. Through Triassic and Early Jurassic time the Nicola Arc is interpreted to have been an extensional arc, with a broad backarc basin (Slide Mountain terrane) situated to the east, with the miogeocline functioning as its continental margin. Diagram B (late Early Jurassic) and Diagram C (early Middle Jurassic) show the beginning of the deformation of the arc and its backarc, with delamination of backarc terranes from their roots and the development of a tectonic wedge of arc rocks being thrust eastward. During the Late Cretaceous and Early Cenozoic, this tectonic wedge developed into a major anticlinorium, the Selkirk Fan, in the Omineca Belt (Price and Monger, 2003).

may have been bordered by an approaching east-facing arc by Early Jurassic time, in front of which developed the Whitehorse Trough, a forearc basin. This is discussed in Chapter 10.

11. Conclusions Within Canada, the extensional, or ‘‘passive’’ western continental margin of Laurentia lasted from the Late Precambrian, between about 600 and 550 Ma, when oceanic crust is presumed to have appeared separating Laurentia from Rodinia, until the Mid-Jurassic, at about 170 Ma. The Antler Orogeny terminated this phase of crustal development considerably earlier in the Great Basin, during the latest Devonian or Early Mississippian (Chapter 11). The duration of the Laurentian Margin in Canada is therefore between 380 and 430 Myr, which is twice the current age of the Atlantic Margin of North America. It is also considerably greater than the duration of the passive-margin phase along any of the other borders of North America. The Arctic Margin of the Franklinian Basin was terminated in northern Ellesmere Island by the Early Silurian (Trettin, 1991). The southern, Ouachita Margin, was closed by collision beginning in the Mid- to Late Mississippian (Chapter 8); that along the Atlantic Margin had an even shorter duration, lasting only until the Middle Ordovician (Chapter 3). Owing to the accretionary tectonism of the Cordilleran Orogeny, the original, thinned crust of the western Laurentian Margin is nowhere preserved. The position of the outer edge of the craton can be established for some

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periods; for example, the Kicking Horse Rim of British Columbia clearly locates the transition from craton to continental slope during much of the Cambrian and Ordovician. However, the deep continental margin basin and the thinned continental- to transitional crust that would be expected to be located outboard of the Columbia and Robson basins, the Selwyn Basin and the Prophet Trough, have long since been deformed, metamorphosed, displaced laterally by transcurrent faulting, and buried beneath the terranes accreted since the Jurassic. Interpretation of Lithoprobe data (Cook et al., 1995) suggests that in southern British Columbia thinned continental crust extends beneath the accreted terranes of the Omineca and Intermontane belts for some 500 km west of the Kicking Horse Rim.

ACKNOWLEDGMENTS The author is very grateful to Barny Poole, Elizabeth Miller and David Morrow for their many valuable comments and suggestions.

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A., 1986, Detachment faulting and the evolution of passive continental margins. Geology, v. 14, pp. 246–250. Lister, G. S., Etheridge, M. A., and Symonds, P. A., 1991, Detachment models for the formation of passive continental margins. Tectonics, v. 10, pp. 1038–1064. McIlreath, I. A., 1977, Accumulation of a Middle Cambrian, deepwater limestone debris apron adjacent to a vertical, submarine carbonate escarpment, southern Rocky Mountains, Canada, in Cook, H. E. and Enos, P. eds., Deep-water carbonate environments, Society of Economic Paleontologists and Mineralogists, Special Publication, v. 25, pp. 113–124. McKenzie, D. P., 1978, Some remarks on the development of sedimentary basins. Earth and Planetary Science Letters, v. 40, pp. 25–32. Miller, E. L., Miller, M. M., Stevens, C. H., Wright, J. E., and Madrid, R., 1992, Late Paleozoic paleogeographic and tectonic evolution of the western Cordillera, in Burchﬁel, B. C., Lipman, P. W., and Zoback, M. L. eds., The Cordilleran Orogen. Conterminous U.S., Geological Society of America, The Geology of North America, v. G-3, pp. 57–106. Monger, J. W. H., and Price, R. A., 1979, Geodynamic evolution of the Canadian Cordillera — Progress and problems. Canadian Journal of Earth Sciences, v. 16, pp. 770–791. Mossop, G. D., and Shetsen, I., compilers, 1994, Geological atlas of the western Canada Sedimentary Basin. Canadian Society of Petroleum Geologists, 510 pp. North American Commission on Stratigraphic Nomenclature, 1983, North American stratigraphic code. American Association of Petroleum Geologists Bulletin, v. 67, pp. 841–875. O’Connell, S. C., 1994, Geological history of the Peace River Arch, in Mossop, G. D. and Shetsen, I., compilers, Geological Atlas of the Western Canada Sedimentary Basin. Canadian Society of Petroleum Geologists, Chapter 28. Oliver, J., 1982, Tracing surface features to great depths: a powerful means for exploring the deep crust. Tectonophysics, v. 81, pp. 257–272. Palmer, A. R., 1960, Some aspects of the Early Upper Cambrian stratigraphy of White Pine County, Nevada and vicinity, in Intermountain Association Petroleum Geologists, Guidebook to the Geology of east central Nevada, pp. 53–58. Poole, F. G., Sandberg, C. A., and Boucot, A. J., 1977, Silurian and Devonian paleogeography of the western United States, in Stewart, J. H., Stevens, C. H., and Fritsche, A. E. eds., Paleozoic paleogeography of the western United States. Society of Economic Paleontologists and Mineralogists, Paciﬁc Section, Paciﬁc Coast Paleogeography Symposium 1, pp. 39–65. Poole, F. G., Stewart, J. H., Palmer, A. R., Sandberg, C. A., Madrid, R. J., Ross, R. J., Jr., Hintze, L. F., Miller, M. M., and Wrucke, C. T., 1992, Latest Precambrian to latest Devonian time; development of a continental margin, in Burchﬁel, B. C., Lipman, P. W., and Zoback, M. L. eds., The Cordilleran Orogen. Conterminous U.S., Geological Society of America, The Geology of North America, v. G-3, pp. 9–56. Price, R. A., and Mountjoy, E. W., 1970, Geologic structure of the Canadian Rocky Mountains between Bow and Athabasca Rivers — a progress report, Geological Association of Canada Special Paper 6, pp. 7–26. Price, R. A., and Monger, J. W. H., 2003, A transect of the southern Canadian Cordillera from Calgary to Vancouver, a ﬁeld trip guidebook, Vancouver, Geological Association of Canada, Cordilleran Section, 165 pp. Rich, M., 1977, Pennsylvanian paleogeographic patterns in the western United States, in Stewart, J. H., Stevens, C. H., and Fritsche, A. E. eds., Paleozoic paleogeography of the western United States: Society of Economic Paleontologists and Mineralogists, Paciﬁc Section, Paciﬁc Coast Paleogeography Symposium, pp. 87–111. Richards, B. C., 1989, Upper Kaskaskia Sequence: uppermost Devonian and Lower Carboniferous, in Ricketts, B. D. ed., Western Canada Sedimentary Basin, Canadian Society of Petroleum Geologists, pp. 165–201. Ricketts, B. D., ed., 1989, Western Canada Sedimentary Basin, Canadian Society of Petroleum Geologists, 320 pp. Robison, R. A., 1960, Lower and Middle Cambrian stratigraphy of the eastern Great Basin, Intermountain Association of Petroleum Geologists, Guidebook to the Geology of East Central Nevada, pp. 43–52. Saltus, R. W., and Hudson, T. L., 2007, Regional magnetic anomalies, crustal strength, and the location of the northern Cordilleran foldand-thrust belt. Geology, v. 35, pp. 567–570. Sloss, L. L., 1963, Sequences in the cratonic interior of North America. Geological Society of America Bulletin, v. 74, pp. 93–113. Sloss, L. L., 1988, Tectonic evolution of the craton in Phanerozoic time, in Sloss, L. L. ed., Sedimentary cover — North American Craton: U.S. Boulder, Colorado, Geological Society of America, The Geology of North America, v. D-2, pp. 25–51. Surlyk, F., Clemmensen, L. B., and Larsen, H. C., 1981, Post-Paleozoic evolution of the East Greenland continental margin, in Kerr, J. 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Trettin, H. P., ed., 1991, Geology of the Innuitian Orogen and Arctic Platform of Canada and Greenland. Ottawa: Geological Survey of Canada, Geology of Canada, v. 3, 569 pp. Walcott, C. D., 1928, Pre-Devonian sedimentation in southern Canadian Rocky Mountains. Smithsonian Miscellaneous Collections, v. 75(4), pp. 147–173. Watts, A. B., 1981, The U. S. Atlantic margin: subsidence history, crustal structure and thermal evolution. American Association of Petroleum Geologists, Education Course Notes Series #19, Chap. 2, 75 pp. Watts, A. B., 1989, Lithospheric ﬂexure due to prograding sediment loads: implications for the origin of ofﬂap/onlap patterns in sedimentary basins. Basin Research, v. 2, pp. 133–144. Watts, A. B., Karner, G. D., and Steckler, M. S., 1982, Lithospheric ﬂexure and the evolution of sedimentary basins, in Kent, P., Bott, M. H. P., McKenzie, D. P., and Williams, C. A. eds., The evolution of sedimentary basins, Philosophical Transactions of the Royal Society, London, A305, pp. 249–281. Welford, K. J., Clowes, R. M., Ellis, R. M., Spence, G. D., Asudeh, I., and Hajnal, Z., 2001, Lithospheric structure across the craton– Cordilleran transition of northeastern British Columbia. Canadian Journal of Earth Sciences, v. 38, pp. 1169–1189.

CHAPTER 6

The Maritimes Basin of Atlantic Canada: Basin Creation and Destruction in the Collisional Zone of Pangea Martin R. Gibling, N. Culshaw, M.C. Rygel and V. Pascucci

Contents 1. Introduction 2. Importance for History of Geology and Sedimentary Research 3. Basement Rocks and Basinal Overview 3.1. Terrane assembly 3.2. Stratigraphic terminology 3.3. Overview of basinal events 3.4. Gondwanan glaciation 4. Mid- to Late Devonian: End of the Acadian Orogeny and Development of Local Extensional Basins 4.1. Termination of the Acadian Orogeny 4.2. Basins and batholiths 5. Late Devonian–Mississippian: A Regional Suite of Extensional Basins 5.1. Horton Group 5.2. Fountain Lake Group 6. Mississippian: Tectonism and Local Basin Filling 7. Mississippian: Global Transgression and Thermal Subsidence 7.1. Windsor Group 7.2. Mabou Group and basin-margin facies 8. Mississippian–Pennsylvanian Unconformity: Onset of a Major Phase of Gondwanan Glaciation 9. Mid-Carboniferous: Extensional Basins and First Phase of Coal Measures 9.1. Lower part of the Cumberland Group 9.2. Stellarton Group 9.3. Southwestern maritimes basin 9.4. Salt migration 10. Pennsylvanian-Permian: Thermally Subsiding Basins and Second Phase of Coal Measures 10.1. Thermal subsidence 10.2. Morien and Pictou groups 11. Pennsylvanian to Permian Sedimentation and Tectonic Events: Late Stages of Pangean Assembly 11.1. Permian eolian sandstones 11.2. Final tectonic events 11.3. Climate change 12. Permian to Mesozoic: End of Maritimes Basin Deposition, and the Breakup of Pangea 13. Synopsis of Maritimes Basin History 13.1. Key points 13.2. Magmatic history 14. A Modern Analogue: Turkey and Eastern Mediterranean Acknowledgments References

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Abstract During the final assembly of Pangea, the Maritimes Basin of Atlantic Canada was tectonically active for B120 Myr from the Mid-Devonian to the Early Permian, following terrane accretion and ocean closure in the region. The basin’s history records a prolonged period of convergence that post-dated the collision of Gondwana and Laurussia. The 12 km of basin Sedimentary Basins of the World, Volume 5 ISSN 1874-5997, DOI 10.1016/S1874-5997(08)00006-3

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fill was laid down in suites of periodically connected depocenters, and parts of the region experienced a polycyclic basin history, with repeated subsidence and inversion of fault-bounded depocenters, many associated with strike-slip faults. During two periods in the basin history, sedimentation overstepped fault zones under a regime of thermal subsidence to blanket much of the region. The basin fills are largely continental but include one open-marine interval with evaporite accumulation (Mississippian), as well as restricted-marine intervals, reflecting progressive loss of oceanic connection. Basinal architecture testifies to rapid subsidence against a backdrop of glacioeustatic influence in a paleoequatorial setting. Volcanics and intrusions were especially prominent during Devonian to Mississippian convergence, and halokinesis greatly influenced later basin development. A partial analogue for the Maritimes Basin is provided by modern Turkey and environs, situated in the Arabia–Eurasia collision zone, where strike-slip faults and basin formation record continued post-collisional convergence adjacent to the Zagros thrust belt. Local crustal thickening, delamination of subducting crust, volcanism, extensional zones, and basin creation along crustal-scale faults are prominent in this region.

1. Introduction The Maritimes Basin comprises a complex suite of Upper Paleozoic strata that covers a large onshore and offshore area of eastern Canada. As presently preserved, the basin extends about 1,700 km from southwestern New Brunswick to the continental margin on the eastern Grand Banks, and about 1,000 km from the southern Grand Banks to offshore Labrador (Figure 1). The term ‘‘Maritimes Basin’’ has emerged as a collective term for a complex stratal package up to 12 km thick that formed over a period of about 120 Myr within an active tectonic framework. Earlier workers recognized a series of variably connected or isolated depocenters (commonly

Figure 1 The Maritimes Basin of Atlantic Canada. Place names referred to in text: A, Antigonish; C, Cobequid Highlands; CA, Caledonia Highlands; CB, Cape Breton Island (Nova Scotia); H, Horton Blu¡; J, Joggins; LAB, Labrador; M, Miguasha; MI, Magdalen Islands (Que¤bec); NF, Newfoundland; NB, New Brunswick; NS, mainland Nova Scotia; QUE, Que¤bec. Fault Zones: CCFZ, Cobequid --- Chedabucto Fault Zone; CFZ, Cabot Fault Zone; HFZ, Hollow Fault Zone. Basins (component depocenters of Maritimes Basin; circled): A, St. Anthony Basin; B, Bay St. George Basin; C, Cumberland Basin; CCP, Central Carboniferous Platform of New Brunswick; D, Deer Lake Basin; G, Gulf of St. Lawrence Basin; M, Moncton Basin; S, Sydney Basin; SB, Stellarton Basin; SM, St. Mary’s Basin; SMB, general area of South Mountain Batholith. Modi¢ed from Bell and Howie (1990).

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considered ‘‘basins’’ in their own right: Figure 1), now fragmented or juxtaposed by contemporaneous and later deformation. Positioned within the collisional zone between the landmasses of Laurussia and Gondwana, the Maritimes Basin records the closure of the Rheic Ocean in an equatorial latitude and the last stages of lithospheric convergence that created the supercontinent of Pangea. The assembly of this part of Pangea represents the first application of plate tectonics to the geological record (Wilson, 1966). Oceanic crust in the region was consumed during the closure of the Iapetus Ocean in the Early Paleozoic, with a suture zone that runs through eastern Canada, and the Maritimes Basin strata rest on continental crust of varying composition and age. Given this dynamic tectonic setting, the complexity of the basin fill is not surprising. At least six basins of varied style developed successively in various parts of the region, and the fills of several of them are superimposed locally, resulting in successive phases of basin subsidence and inversion, reactivation of tectonic lineaments, halokinesis, and erosion of earlier strata (St. Peter, 1993). This polycyclic basinal history has important consequences for thermal maturation, diagenesis, and the migration of hydrocarbons and mineralizing fluids. The restless plates were in motion again by the Late Permian, reactivating Upper Paleozoic lineaments as Pangea fragmented and the Atlantic Ocean began to open. The present overview builds on comprehensive regional accounts (Bell and Howie, 1990; van de Poll et al., 1995), as well as local studies (particularly Knight, 1983; MacLean and Wade, 1992; St. Peter, 1993; Rehill, 1996; Calder, 1998; Giles and Utting, 1999; Pascucci et al., 2000). We emphasize the stratigraphy and detailed stratal architecture of three of the larger depocenters — the Cumberland, Gulf of St. Lawrence and Sydney basins (Figure 1) — and conclude by suggesting that Turkey and the eastern Mediterranean, situated between the converging plates of Eurasia and Arabia, provides a partial modern analogue for the Maritimes Basin.

2. Importance for History of Geology and Sedimentary Research Upper Paleozoic outcrops of the Maritimes Basin came to world attention in the 1840s and subsequent decades, when Sir Charles Lyell, Sir William Logan, and Sir John Dawson described the remarkable seacliff sections and their fossils (Scott, 1998; Rygel and Shipley, 2005; Falcon-Lang and Calder, 2005; Calder, 2006), culminating in Dawson’s comprehensive account of the region in Acadian Geology (1855). The hundreds of kilometers of coastal cliffs have resulted in the most complete exposures of Pennsylvanian coal measures anywhere in the world. In particular, the cliffs at Joggins with their splendid exposures of coals and standing trees were to Lyell what the Galapagos Islands were to Charles Darwin, as he unraveled the evolution of landscapes through time (Calder, 2006). He subsequently considered Joggins to represent the world’s best exposure of Carboniferous coal measures (Lyell, 1871). Joggins was mentioned (although not named as such) by Darwin in Origin of Species, featured prominently in the evolution debate of the late 1850s (Wilberforce, 1860), and is now a UNESCO World Heritage site as the world’s iconic coal-measure section. Joggins was awarded UNESCO status in the summer of 2008. Additionally, the Maritimes Basin contains the Upper Devonian UNESCO World Heritage site at Miguasha in southern Que´bec, which has yielded material for exhibits and teaching collections around the world. First found by Abraham Gesner in 1842, the locality was rediscovered in 1879, and its remarkable fish faunas have been influential in understanding the origins of tetrapods, ray-finned fishes and coelacanths, and the transition from aquatic to terrestrial vertebrates (Schultze and Cloutier, 1996). The Maritimes Basin also contains some of the world’s best preserved Devonian, Carboniferous, and Permian forested levels and root systems (Calder et al., 1996; Elick et al., 1998; Calder et al., 2006; Falcon-Lang, 2006), remarkable instances of sediment–vegetation interaction (Rygel et al., 2004, 2006), and the earliest known reptile and land snail (at Joggins), entombed inside tree trunks or in seasonal waterholes (Lyell and Dawson, 1853; Falcon-Lang et al., 2004). The basin is also the only known area where the tetrapod trackway record can be documented from the Devonian to the Permian (Hunt et al., 2004), and strata at Horton Bluff in Nova Scotia have yielded the world’s most complete assemblage of Tournaisian tetrapod material and trackways, providing key vertebrate information for the interval referred to as ‘‘Romer’s Gap’’. The basin also contains the first evidence for gregarious behavior in earth history, as groups of tetrapods moved around the only known forest of the early conifer Walchia (Van Allen et al., 2005). The exceptional coastal exposures have allowed for detailed analysis of sequence development under the icehouse conditions of the Upper Paleozoic (e.g., Davies and Gibling, 2003; Giles and Boutilier, 2003; Gibling et al., 2004). The presence of rapidly subsiding extensional basins and regions affected by slow thermal subsidence has permitted the investigation of tectonic, glacioeustatic, and climatic factors in high- and low-accommodation settings, and the excellent floral record has allowed investigation of the influence of the developing lowland and upland vegetation cover and wildfire events on landscape evolution (Falcon-Lang, 2000, 2003a, 2003b, 2004; Falcon-Lang and Bashforth, 2004, 2005). A thick Mississippian evaporite succession with diapir fields, parts of which are exposed in the cliffs, provides considerable insight into the effects of halokinesis on basin development

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and sedimentation (Alsop et al., 2000; Waldron and Rygel, 2005). From an economic viewpoint, Abraham Gesner conducted the first extraction of kerosene in 1836 (Brice, 2002), major coalfields formerly supplied most of Canada’s needs, hydrocarbons are produced locally, and the region currently produces large amounts of gypsum and rock salt.

3. Basement Rocks and Basinal Overview 3.1. Terrane assembly Following the breakup of Rodinia in the Neoproterozoic, the southern edge of Laurentia was characterized by an irregular margin with salients and reentrants (Figure 8 in Chapter 1; Lavoie, Chapter 3, this volume), and alongstrike variations in the margin’s geological history during the Paleozoic reflect this irregularity. Much of the Maritimes Basin lies within the Que´bec Reentrant, between the New York and St. Lawrence promontories, and this position may account for the roughly elliptical form of the preserved basinal strata (Figure 1). By the Mid-Devonian, Atlantic Canada consisted of a collage of diverse terranes (Figure 2) assembled during Late Precambrian to Early Paleozoic orogenic events that included the Taconian, Salinian, and Acadian orogenies (Marillier et al., 1989; Barr et al., 1998; Van Staal et al., 1998; Lavoie, Chapter 3, this volume). The Proterozoic Grenville Province of the Canadian Shield lay along the margin of Laurussia in southern Que´bec, where a Lower Paleozoic carbonate platform developed (Figure 1). The Humber Terrane and the Blair River Inlier represent these rocks in Atlantic Canada (Figure 2). West Avalonia — of peri-Gondwanan affinity — had rifted from Gondwana and transferred to the North American margin by the Mid-Silurian, closing the Iapetus Ocean in front along the Iapetus Suture (Figure 3) and opening the Rheic Ocean behind (Nance et al., 2002). West Avalonia is represented by the Avalon and Mira terranes, and other terranes were emplaced between the converging continental blocks (Figure 2). The Early to Mid-Devonian Acadian Orogeny has commonly been attributed to the docking of the periGondwanan Meguma Terrane that underlies southern Nova Scotia, although the Avalon and Meguma terranes may have amalgamated during the Lower Paleozoic (Murphy et al., 2004). Keen et al. (1991) noted northdipping reflectors below the Gulf of Maine, and suggested that they represent subducted oceanic lithosphere, perhaps connected with the Acadian Orogeny. Continued convergence following these collisional events represents the movement of the West African segment of Gondwana (Gondwana itself is not represented in Atlantic Canada) against the North American margin (Figure 3). Although the Maritimes Basin can in some respects be considered a ‘‘successor basin’’ following the Acadian Orogeny, this term fails to represent adequately

Figure 2 (1998).

Terranes in the northern Appalachian Orogen. See Figure 1 inset for general location. Modi¢ed from van Staal et al.

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Figure 3 Tectonic setting of Maritimes Basin within collisional area of Gondwana and Laurussia in the ¢nal stages of Pangean assembly. Modi¢ed from Calder (1998).

the intricate interplay of tectonics and sedimentation that continued until Pangean assembly was complete. The basin fill broadly constitutes an ‘‘overlap succession’’ that links all the tectonostratigraphic basement terranes in the northern Appalachian orogen (White and Barr, 1998). Paleogeographic maps for the Devonian to Permian show the progressive loss of oceanic connections in eastern Canada (Torsvik and Cocks, 2004; see Figures 10–13 in Chapter 1), until the marine realm was represented only by restricted-marine incursions from the Mid-Euramerican Sea during the Pennsylvanian (Figure 3). A narrow seaway may have been re-established briefly during Early Permian oblique rifting, prior to the final assembly of Pangea by the Mid- to Late Permian (Vai, 2003).

3.2. Stratigraphic terminology Because a large proportion of the basinal strata lie under the sea (Figure 1), the stratigraphy of inherently complex strata at the basin margins has received the most attention, but we emphasize here more recent offshore studies

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Figure 4 Stratal groups widely recognized within the Maritimes Basin ¢ll. Names in parentheses are local equivalents.

that have provided a fuller account of the basin as a whole. Inevitably, this approach tends to oversimplify the basin history. Recent stratigraphic useage has defined a series of groups that can be recognized across much of Atlantic Canada (Figure 4), allowing local complexities to be resolved at the formation and member level. The groups have a restricted age range, and represent a series of local to near-basinwide events. The group approach tends to break down where small, localized depocenters existed and where fault activity within larger depocenters generated local basin-margin facies. Additionally, in the case of the Cumberland and Pictou groups, group identification is based on the prevailing red or gray coloration of the strata and the presence or absence of coal (Figure 4; Ryan et al., 1991); these attributes primarily reflect the spatially variable factor of groundwater level, and the ‘‘groups’’ are lateral equivalents in places. On account of the scattered distribution and poor exposure of Devonian parts of the basin fill, no group designation has yet been applied. A capping succession of largely eolian redbeds in the Gulf of St. Lawrence is currently unnamed.

3.3. Overview of basinal events The basin’s Mid-Devonian to Permian history (Figure 5) commenced with a series of Devonian basinal events (1) followed or accompanied by the intrusion of the South Mountain Batholith (2), after which two Late Devonian to Mississippian extensional basin successions developed in various parts of the region (3, 4). A widespread Mississippian marine to lacustrine unit (5) represents a period of thermally driven subsidence and Visean transgression. Volcanic activity took place locally during these episodes. The Mississippian — Pennsylvanian Unconformity, probably linked to the onset of a major phase of glaciation, is well represented in the region and may have been enhanced by the onset of Alleghanian tectonism. Renewed tectonism resulted in

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Figure 5 General tectono-stratigraphic overview of the Maritimes Basin, with seven major basinal phases and intervening periods of inversion during B130 Myr. Two unconformities have received regional names (Giles and Utting, 1999).

a series of fault-related basins (6), after which a period of thermally driven subsidence broadened the basin (7), with regional overstep above the Early Westphalian Unconformity. Coals are prominent in events 6 and 7, and diapirism of Mississippian evaporites continued through this period. Unnamed Permian sandstones (8) cap the Maritimes Basin fill, after which Late Permian to Cenozoic sediments accumulated as Pangea moved to rift and drift phases. Two periods of near-basinwide thermal subsidence are apparent in the basin history, and periods of inversion (I) separated local basinal filling periods.

3.4. Gondwanan glaciation The Late Paleozoic ice age spanned the Late Devonian to early Late Permian, when Gondwana lay near the south pole (Caputo and Crowell, 1985). In paleoequatorial areas such as the Maritimes Basin, the effects of glacially driven eustasy have been inferred from cyclic sequences and incision in coastal areas (Gibling and Bird, 1994; Smith and Read, 2000; Wright and Vanstone, 2001), as well as from isotopic analysis of carbonates and organic matter (Mii et al., 2001). However, the timing, extent, and magnitude of glaciation are the subject of considerable debate (Isbell et al., 2003; Jones and Fielding, 2004). In a recent summary of the Gondwanan record, Isbell et al. (2003) documented phases of alpine glaciation in the Frasnian to Tournaisian (Glacial I) and the Namurian to Early Westphalian (Glacial II), followed by the development of ice sheets in the Stephanian to Permian (Glacial III). They inferred that the ice sequestered during Glacials I and II was insufficient to create large eustatic fluctuations and cycle formation in equatorial latitudes. Other researchers, however, have suggested more extensive and prolonged ice coverage (Crowley and Baum, 1991; Gonzalez-Bonorino and Eyles, 1995). Cyclic successions in the Maritimes Basin and elsewhere in equatorial Euramerica have been taken to imply glacioeustatic events in the Late Visean (Asbian and Brigantian: Smith and Read, 2000; Wright and Vanstone, 2001; Giles and Boutilier, 2003) and in the later Westphalian (Gibling et al., 2004). As noted by Gibling and Rygel (in press), the cyclic successions are especially well developed during periods of thermal subsidence, which provides circumstantial support for a glacioeustatic origin. Additionally, in the Maritimes Basin and at many

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localities worldwide, an unconformity accompanied by a major biostratigraphic change lies at the Mississippian– Pennsylvanian boundary, and is widely interpreted to mark a major Gondwanan glacial period (equivalent to Glacial II but seemingly of considerable magnitude). This event was associated with profound sea-level lowering and continent-wide landscape exposure (Saunders and Ramsbottom, 1986; Beuthin, 1994), although the unconformities may be of tectonic origin in places (Ettensohn, 1994).

4. Mid- to Late Devonian: End of the Acadian Orogeny and Development of Local Extensional Basins 4.1. Termination of the Acadian Orogeny This period of the basin’s history has the most fragmentary record, but represents some 40 Myr (Figure 6) — about a third of the basin’s history. Small, poorly exposed, and tectonically disturbed areas of alluvial and lacustrine strata, with thick volcanic successions and some intrusions, are present in Nova Scotia and the Gaspe´ Peninsula of southern Que´bec. The strata may have been laid down in small, fault-bounded basins along reactivated tectonic lineaments, but no coherent model is available to account for their presence. A brief survey of Acadian events across the region, as summarized by Lavoie (Chapter 3, this volume), illustrates the range of deformational styles and the diachronous history of orogeny in the region (Figure 6). Along the coast of western Newfoundland and just offshore, an Acadian thrust front is represented by thinskinned thrusts and a foreland basin succession that includes the Emsian Red Island Road Formation (Stockmal et al., 1998). Subsequently, the Humber Arm Allochthon was emplaced over these rocks during a thick-skinned deformation event, probably during the Mid-Devonian. On the Meguma Terrane of southern mainland Nova Scotia, the Torbrook Formation of the Meguma Group ranges up into the Emsian (Schenk, 1995), and the Acadian Orogeny is represented by low-grade regional metamorphism of the Meguma Group at 395–380 Ma (Middle Devonian) (Hicks et al., 1999). In the Gaspe´ area, a 5 km succession of Early to Mid-Devonian strata reflects oblique collision of West Avalonia with the St. Lawrence Promontory of the Canadian Shield (Lawrence and Williams, 1987; Rust et al., 1989). In the northern United States, the Acadian Orogeny generated the Catskill Delta, which advanced westward into a foreland basin during the Middle and Late Devonian (Faill, 1985; Ettensohn, Chapter 4, this volume). In general, Acadian deformation in Atlantic Canada took place during the Early to Middle Devonian, with somewhat younger ages along the Appalachian chain to the southwest. We include in the Maritimes Basin fill several successions of Mid- and Late Devonian age (possibly as old as latest Early Devonian) that appear to have post-dated major Acadian episodes in their respective areas but may be coeval with or even older than Acadian events elsewhere. Across cratonic areas of western Canada and North America generally, topmost Devonian and

Figure 6 Devonian strata in selected parts of Atlantic Canada. Sources of data: Gaspe¤: Rust et al. (1989), Jutras and Prichonnet (2002); Central New Brunswick: McCutcheon (1990); South Mainland Nova Scotia: Schenk (1995), Clarke et al. (1997), Hicks et al. (1999); North Mainland Nova Scotia: Martel and Gibling (1996), Cormier et al. (1995), Murphy (2001), Pe-Piper and Piper (2002); Cape Breton Island: Barr et al. (1995),White and Barr (1998); Western Newfoundland: Knight (1983), Stockmal et al. (1998). Age dates from Tucker et al. (1998). Note that ages and relationships of many formations have considerable uncertainty.

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Mississippian strata rest upon a major unconformity that marks the base of the Upper Kaskaskia sequence (Sloss, 1963; Richards, 1989). In view of the active collisional setting of Atlantic Canada at this time, we have not attempted to reconcile events in the local stratigraphic succession with those documented from cratonic areas elsewhere.

4.2. Basins and batholiths In central Cape Breton Island, the McAdams Lake Formation comprises nearly 2 km of strata that include oil shales and thin coals (Figure 7), probably of Middle Devonian age (White and Barr, 1998). The only evidence of igneous activity is an undated syenite. The formation was laid down in an extensional basin along a reactivated terrane boundary, and the strata were deformed prior to the Mississippian. The Fisset Brook Formation includes Late Devonian rhyolites (37374 Ma; Barr et al., 1995). On mainland Nova Scotia, the Murphy Brook Formation in the Cobequid Hills has been dated as Emsian to Eifelian based on plant remains (Calder, 1998), although Pe-Piper and Piper (2002) suggested that the age should be reassessed because clast lithologies do not match known available source rocks. The McAras Brook Formation farther east comprises mafic flows and clastics (Fralick and Schenk, 1981; Murphy, 2001). The Sunnyville Formation of the eastern mainland includes more than 1,400 m of strata, largely mafic and felsic flows and pyroclastic rocks, with gabbroic plutons, associated with clastic strata of the Glenkeen Formation and other units (Cormier et al., 1995); a rhyolite sample yielded a U/Pb date of 38972 Ma (Givetian). In central New Brunswick, the Piskahegan Formation, at least 450 m thick, includes rhyolite and some mafic flows, tuffs, ignimbrites, and shallow intrusives that form part of the Mount Pleasant caldera complex (McCutcheon, 1990). On the southern side of the Gaspe´, conglomerates of the Fleurant Formation are overlain by the turbidite facies of the Escuminac Formation, with abundant fossil fish (Hesse and Sawh, 1992). The paleoecology of the fossils and geochemical evidence suggest an estuarine environment (Schultze and Cloutier, 1996). Elsewhere in the Gaspe´, post-Acadian units are present in a small graben (Jutras and Prichonnet, 2002). An interesting aspect of this period is the intrusion of the South Mountain Batholith in the Meguma Terrane south of the Cobequid–Chedabucto Fault Zone. The batholith constitutes the largest granitic body in the northern Appalachians (Clarke et al., 1997), and is familiar to generations of visitors to the Peggy’s Cove lighthouse. This important phase of felsic magmatism is bracketed at 375710 Ma (Late Devonian), with dates of

Figure 7 Stratigraphic column for the Sydney Basin. Modi¢ed from Boehner (1985) and Pascucci et al. (2000). Age dates from Gradstein et al. (2004).

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37273 (Famennian) for the main belt of plutons. Local plutons may range as late as 35771 Ma (Maclean et al., 2003), and mafic intrusions are present locally (Tate and Clarke, 1995). Although there is no direct evidence that the batholith had a volcanic expression, Clarke and Bogutyn (2003) suggested that oscillatory zoning in micas reflects rapid changes in pressure during eruptive phases; however, they noted that rhyolite fragments appear to be absent from younger conglomerates that overlie the batholith. The felsic magmas were emplaced at about 12 km depth and were unroofed by the latest Famennian, prior to 359 Ma, implying very rapid exhumation — in the order of 1 mm/year (Martel and Gibling, 1996; Clarke et al., 1997; Kontak et al., 2004). Despite the proximity of the batholith to coeval basins north of the fault, the relationship (if any) between igneous and basinal events remains obscure and, at the time of intrusion, the South Mountain Batholith may not have lain adjacent to northern Nova Scotia.

5. Late Devonian–Mississippian: A Regional Suite of Extensional Basins 5.1. Horton Group The Late Devonian marked the development of a suite of linear fault-bounded basins (mainly half-grabens) across the region, linked to extension on master faults oriented E-W to NE-SW, largely along the structural grain of the basement. The basin fills are typically several kilometers thick (Figures 7–10) and extend across terrane boundaries, including the Avalon/Meguma boundary (Murphy and Rice, 1998). They are well imaged on seismic profiles across the offshore Sydney and Gulf of St. Lawrence basins (Figures 8 and 10). Below Prince Edward Island and the southern Gulf of St. Lawrence, the Malpeque Basin is unusually thick, containing 8 km of strata, and thinner accumulations are present in sag basins and veneers over older bedrock (Durling and Marillier, 1993). Deep-crustal seismic profiles suggest that the basins were generated by the reactivation of Acadian thrusts in extension, probably creating a master detachment (Marillier et al., 1989), and most basinal fills contain thick dark shales with a near-coincident Tournaisian age, suggesting that a regional period of exceptionally rapid basin subsidence outpaced sediment supply. The basin fills are assigned to the Fountain Lake and Horton groups and their equivalents (Figure 4), and all across the region are typically tripartite with basal alluvial clastics, medial lacustrine and restricted-marine deposits, and a topmost alluvial succession (Hamblin and Rust, 1989; St. Peter, 1993; Murphy et al., 1995; Martel and Gibling, 1996; Murphy and Rice, 1998; Te´nie`re et al., 2005). Basin-margin alluvial fans with coarse conglomerates are prominent in many basins, and some units contain paleoplacer gold eroded from adjacent bedrock. Prior to Horton deposition, a period of intense weathering, probably under tropical conditions, yielded up to 6 m of saprolite on the South Mountain Batholith (O’Beirne-Ryan and Zentilli, 2003). The shale-rich medial intervals comprise stacked shoaling-up cycles or parasequences that suggest repeated increments of subsidence on basin-bounding faults (Martel and Gibling, 1991). Although these beds have traditionally been considered lacustrine, discoveries of calcareous algae, echinoderms, agglutinated foraminifera, and ostracodes at Horton Bluff indicate local marine influence (Mamet, 1995; Tibert and Scott, 1999). Paleogeographic maps suggest that a tongue of the Rheic Ocean extended between Laurussia and Gondwana at this time, close to Atlantic Canada. The strata also exhibit considerable evidence for instability in the form of collapse structures, slumps, and clastic dykes (Knight, 1983; Martel and Gibling, 1993), some of which may have been associated with earthquake activity. Oil shales are especially prominent in the Moncton Basin of

Figure 8 Interpreted seismic pro¢le to illustrate aspects of o¡shore Sydney Basin succession. Line 4101-83, modi¢ed from Pascucci et al. (2000). Location shown in Figure 1.

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Figure 9 Simpli¢ed cross-section through Prince Edward Island and part of the southern Gulf of St. Lawrence Basin (see Figure 6 for general location). Colors within Mabou and Cumberland Group represent locally recognized formations. Modi¢ed from Rehill (1996) and Giles and Utting (1999), with an early version of this diagram courtesy of P.S. Giles. Rehill (1996) assigned the Bradelle and Green Gables formations to the Cumberland Group and overlying formations to the Pictou Group. The term ‘‘Prince Edward Island Group’’ has been used for strata exposed on Prince Edward Island (van de Poll, 1989) but has not been widely adopted.

Figure 10 Interpreted seismic pro¢les under the Gulf of St. Lawrence from Lithoprobe line 86-1. Modi¢ed from Marillier et al. (1989). Locations shown in Figure 1.

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southern New Brunswick, where strata of the Albert Formation have been mined and retorted and contain hydrocarbon reservoirs (Smith and Gibling, 1987; St. Peter, 1993). The Horton Group is considered the main source-rock interval of the Maritimes Basin, and outcrop belts are widely associated with oil seeps (Hamblin et al., 1995). In western Newfoundland, the Anguille Group (Horton equivalent) is present in the Bay St. George and Deer Lake basins, as well as at isolated localities as far north as Conche and the offshore St. Anthony Basin (Figure 1; Jansa and Mamet, 1984; Bell and Howie, 1990; Hamblin et al., 1995). Basin formation was associated with the Cabot Fault Zone, where positive flower structures suggest strike-slip activity (Hyde et al., 1988; Wright et al., 1996). Coarse clastic facies of the Horton Group are also widespread across the southern Grand Banks, where they are truncated by erosion at the continental margin (Bell and Howie, 1990). Topmost Anguille strata onlap basement rocks in the Bay St. George Basin in southwestern Newfoundland (Hall et al., 1992), as do Horton strata beneath the Gulf of St. Lawrence (Durling and Marillier, 1993), implying that a period of thermal subsidence succeeded the main period of fault-related subsidence.

5.2. Fountain Lake Group In the Cobequid Hills of mainland Nova Scotia, the Fountain Lake Group includes several kilometers of rhyolite and basalt, associated with granitic, gabbroic, and dioritic intrusions. Earlier workers considered that these rocks predated the Horton Group, but recent U/Pb dating has yielded Famennian to Tournaisian (362–350 Ma) dates (Pe-Piper and Piper, 2002). The volcanics may have accumulated in a very rapid eruptive phase, perhaps spanning as little as 3–10 Ma (Dessureau et al., 2000). Pe-Piper and Piper (2002) suggested that volcanism was linked to the Cobequid Shear Zone, which is known to have been active at the time of igneous activity (Pe-Piper et al., 2004). Partial melting under hydrous conditions may have been triggered by ingress of water from lakes down this crustal-scale fault system, which would also have acted as conduits for eruption. Horton strata (Nuttby Formation) in the Cobequid Hills may have been juxtaposed with rocks of the Fountain Lake Group along the Cobequid Shear Zone (Pe-Piper and Piper, 2002).

6. Mississippian: Tectonism and Local Basin Filling Evidence has been accumulating for a Late Tournaisian to Early Visean tectonic and basinal event in parts of Atlantic Canada. The Horton/Windsor group contact is an angular unconformity at numerous localities, including parts of the Gulf of St. Lawrence (Durling and Marillier, 1993), the St. Mary’s Basin of Nova Scotia (Murphy et al., 1995), the Moncton Basin of southern New Brunswick (St. Peter, 1993), and parts of western Cape Breton. The topmost Tournaisian Wilkie Brook Formation near Antigonish, composed of alluvial-fan and playa lake deposits, rests unconformably on the Horton Group but is overlain concordantly by the basal Windsor Group (Boehner and Giles, 1993). In eastern mainland Nova Scotia, Reynolds et al. (2004) documented a phase of deformation and low-grade regional metamorphism, dated at 335–340 Ma, that affected the Horton Group but not the overlying Windsor Group; this tectonic activity may have been linked to displacement and overthrusting along the Avalon-Meguma boundary. In the Cobequid Hills, Piper and Pe-Piper (2001) noted that the unconformity-based Falls Formation of the topmost Horton Group lies close to north-vergent thrusts and mylonites, probably of Late Tournaisian age, and commented that half-graben formation was locally synchronous with pluton emplacement and thrusting. These observations imply a distinct period of tectonism and basin inversion, perhaps along major tectonic lineaments. As noted by Calder (1998), the base of the Windsor Group is Mid-Visean, and a hiatus appears to exist across the region. In southern New Brunswick, an unconformity-bounded, Late Tournaisian basinal fill — formerly included within the Horton Group — has recently been designated as the Sussex Group (Park and St. Peter, 2005). These strata comprise the conglomeratic Round Hill Formation, the shale-rich Weldon Formation, the Gautreau Formation composed of nearly 500 m of shale, limestone, halite, and minor anhydrite and glauberite, and the Boyd Creek Tuff, up to 30 m thick. The glauberite-rich evaporites are considered non-marine.

7. Mississippian: Global Transgression and Thermal Subsidence 7.1. Windsor Group The Mississippian Windsor and Mabou groups mark a widespread change in sedimentation across the region. The Windsor Group, of Mid- to Late Visean age, is up to about 1 km thick in most areas (Figure 11, but is

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Figure 11 Stratigraphy of the Windsor Group in the Shubenacadie Basin of Nova Scotia, based on drill hole SB-1, where strata are well bedded and not a¡ected by salt migration. Modi¢ed from Giles (1981), Boehner (1984) and Giles and Boutillier (2003).

considerably thicker in some areas of active faulting, such as western Newfoundland and southern New Brunswick. The group comprises the only open-marine unit in the basin succession. It underlies much of the present basin area, from southwestern New Brunswick northward to the northern Gulf of St. Lawrence, across the southern Grand Banks (Figure 12), and as far north as the St. Anthony Basin north of Newfoundland (Jansa and Mamet, 1984) (Figure 1). Although Sanford and Grant (1990) mapped the Windsor and Mabou groups within the St. Lawrence estuary, recent seismic studies have not confirmed the presence of Carboniferous strata there (Dietrich et al., 2005). Across the Central Carboniferous Platform in New Brunswick, Windsor carbonates are present on bedrock of the Piskahegan Group in places, although their extent is uncertain (McLeod, 2000). Although there is much local variation, a broadly similar stratigraphy characterizes most outcrop belts. Local coarse conglomerates imply continued tectonic activity in some areas, where attribution to Horton or Windsor groups may be difficult. Marine carbonates of the basal Windsor Group overstep the basin boundaries onto basement rocks across the region, and constitute a distinct seismic reflector package (Figures 8 and 10). Schenk et al. (1994) recorded relief on the basal Windsor unconformity of 600 m over distances of 2,500 m. The group commences with a laminated bituminous limestone and shale (the Macumber Formation and equivalents, typically only a few meters thick: Figures 7 and 11) laid down in relatively deep-water conditions on a quiet, reducing basin floor over an area of at least 250,000 km2. This stratigraphic situation may imply near-simultaneous marine invasion across an area below sea-level (Geldsetzer, 1977; Schenk et al., 1994), although superimposed beach deposits at some localities suggest a more gradual transgression (Jutras et al., 2006). Inundation of higher bedrock areas resulted in a suite of bioherms of the Gays River Formation, up to 60 m thick, composed of crystalline calcite and peloidal micrite with bryozoan thrombolite, cryptomicrobial, and

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Figure 12 Interpreted seismic pro¢le across edge of Burin Platform, south of Newfoundland. Modi¢ed from Pascucci et al. (1999). Location shown in Figure 1.

skeletal facies (Giles et al., 1979; Schenk et al., 1994). The bioherms locally host Pb-Zn deposits. In eastern Cape Breton and in western Newfoundland, the basal Windsor strata include carbonate mounds with tubeworms, developed over fault-related hydrothermal vents (von Bitter et al., 1990). A suite of features that include slumps, vugs, pockmarks, and mud volcanoes reflect downslope movement and high pore-fluid pressure (Knight, 1983; Schenk et al., 2001). Overlying Lower Windsor strata (major cycle 1) include hundreds of meters of sulfate and chloride evaporites, with local potash and borate salts (Figures 7 and 11). The evaporite-rich nature of the Windsor — in contrast to the predominantly carbonate-rich coeval successions of continental United States and Western Europe — indicates the presence of arid embayments and progressive restriction of oceanic connection as Pangean assembly continued. Middle and Upper Windsor strata (major cycles 2–5) comprise cycles with thin, fossiliferous marine carbonates that can be traced individually across much of the region, where they are intercalated with redbeds. Spectral analysis by Giles and Boutilier (2003) suggests that these cycles correspond to Milankovitch frequencies, implying a strong orbital control over relative sea-level fluctuations during Asbian and Brigantian time. Breccia units several hundred meters thick at some localities reflect salt dissolution and collapse (Brisebois, 1981; Boehner, 1986), and some may represent intra-Carboniferous solution events (Boehner, 1985). From southern Gaspe´ southwards through New Brunswick, Windsor equivalents include detrital deposits and thick groundwater calcretes, recently placed within the newly defined Perce´ Group and probably interfingering with evaporitic, marine Windsor deposits to the south (Jutras and Prichonnet, 2002, 2005; Jutras et al., 2007). Volcanic and intrusive rocks of confirmed or probable equivalent age are present on the Magdalen Islands in the Gulf of St. Lawrence, in southern Cape Breton Island, in the Cobequid Hills, and locally on the Central Carboniferous Platform of New Brunswick (Brisebois, 1981; Barr et al., 1985, 1994; Pe-Piper and Piper, 2002; New Brunswick Department of Natural Resources, 2000). The basal Windsor overstep probably reflects in part a global Visean transgression — one of the largest such events in Phanerozoic history (Hallam, 1984). Additionally, Sydney Basin seismic profiles (Figure 8) show that the Windsor Group occupies ‘‘hollows’’ within Mississippian half-grabens. This implies that the strata accumulated as extension-driven subsidence waned and a phase of thermal subsidence set in (Pascucci et al., 2000) and, under such a low-accommodation regime, glacioeustatic events may have generated the prominent cyclic patterns evident in Middle and Upper Windsor strata (Gibling and Rygel, in press). Onlap in the Deer Lake Basin was also attributed to thermal subsidence by Hyde et al. (1988), and a thermal subsidence phase probably affected much of the region. Active tectonism during Windsor Group deposition may have occurred in the Cumberland Basin, where marked changes in thickness (possibly to 3 km locally) occur within fault-bounded depocenters (Waldron and Rygel, 2005). Windsor gypsum, salt, and potash form the basis of important industries, including salt mines within diapirs (Carter, 1987). In the Sydney Basin, evaporitic brines within modern deep groundwater systems are probably the remnants of recharged Windsor brines (Gibling et al., 2000; Martel et al., 2001) that may have contributed to mineralization in the region (Ravenhurst et al., 1989).

7.2. Mabou Group and basin-margin facies The Mabou Group of Upper Visean to Lower Namurian age, up to about 1 km thick, rests conformably on the Windsor Group and its wide distribution suggests the continuance of thermally driven subsidence. Mabou strata

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cover much of the Central Carboniferous Platform in New Brunswick, resting mainly on pre-Carboniferous strata (New Brunswick Department of Natural Resources, 2000). Hamblin (2001) considered these strata to represent the final fill of the rift system within large distensive fault-bounded areas. A widespread transition is observed from gray, shallow lacustrine facies (Hastings Formation) with lakes up to 200 km long and more than 100 km wide, to red playa and floodplain deposits (Pomquet Formation), with calcretes and vertisols implying a relatively arid climate (Hamblin, 2001). The Pomquet Formation locally oversteps the Hastings Formation to rest on Windsor and basement rocks (Boehner and Giles, 1993). Although the strata have been considered nonmarine, Copeland (1957) and Calder (1998) presented faunal evidence, especially from ostracode taxa, for marine influence at some levels. The reduction in marine connection may reflect the onset of a major phase of Gondwanan glaciation and sea-level lowering, or could represent tectonic factors. Although tectonic activity was generally minor, local conglomerate wedges in the Windsor Group imply continued fault motion (Fralick and Schenk, 1981; Knight, 1983; Bradley and Bradley, 1986; Stevens et al., 1999). In the Mabou Group of southern New Brunswick, thick alluvial-fan conglomerates (Hopewell Cape Formation) and basinal equivalents, up to 2 km thick, indicate tectonic activity along the Harvey-Hopewell fault, with local onlap of Mabou strata onto older Carboniferous and basement rocks (St. Peter, 1993; C. St. Peter, personal communication). Formation of Mabou Group minibasins within the Cumberland Basin records localized sediment accumulation in association with halokinesis (Waldron and Rygel, 2005), and these deposits bear striking similarity to the Mabou Group half-grabens described in Cape Breton Island (Hamblin, 2001).

8. Mississippian–Pennsylvanian Unconformity: Onset of a Major Phase of Gondwanan Glaciation A prominent unconformity caps the Mabou Group and cuts down into older strata at numerous localities. The unconformity is particularly well seen in the Sydney Basin (Figures 7 and 8), where seismic profiles indicate a regional erosional surface overlain by Duckmantian coal measures of the Morien Group (Pascucci et al., 2000). The surface is an angular unconformity with a few degrees of discordance, and most faults in the underlying Mississippian strata do not penetrate up into the overlying coal measures. Although the unconformity surface is broadly planar, overlying fluvial strata fill a series of broad, linear valleys that follow the structural grain of the underlying basement rocks. The unconformity marks the most pronounced floral break in the Maritimes Carboniferous (Bell, 1944; Calder, 1998). Beneath Prince Edward Island and the Gulf of St. Lawrence, Rehill (1996) and Giles and Utting (1999) identified the Mississippian–Pennsylvanian Unconformity beneath the Namurian Cumberland Group (Figure 9). Elsewhere in the region, the unconformity is cryptic, but it has been identified beneath the Boss Point Formation (the basal Namurian to Langsettian unit of the Cumberland Group; Figure 13) in the Cumberland and Moncton basins (C. St. Peter, personal communication). The Boss Point Formation marks a fundamental change in the depositional system of the region, with alluvium covering progressively wider areas (Calder, 1998). A comparable unconformity beneath Pennsylvanian coal measures is also present in the Narragansett Basin of New England (Thompson and Hermes, 2003). As noted earlier, unconformities at many localities worldwide probably represent the onset of a major glacial period, with regional lowering of base-level and widespread subaerial erosion (Beuthin, 1994), although tectonism associated with the onset of the Alleghanian Orogeny may have created unconformities at this level in the Appalachian Basin (Ettensohn, 1994). In the Sydney Basin, the angular nature of the unconformity suggests that tectonic activity coincided with base-level change; the relatively young age of strata (Duckmantian– Bolsovian) that rest on the unconformity suggests that the surface represents more than one erosional period. Seismic lines from the Cumberland Basin reveal an angular unconformity between the Boss Point Formation and the Mabou Group, although the angular relationship is not visible in outcrop; Waldron and Rygel (2005) attributed this feature to salt tectonism in the latest Mississippian to earliest Pennsylvanian.

9. Mid-Carboniferous: Extensional Basins and First Phase of Coal Measures 9.1. Lower part of the Cumberland Group The basal strata of the Cumberland Group occupy basins along major fault zones, especially strands of the Cabot, Hollow, and Cobequid-Chedabucto Fault Zones, where subsidence and accumulation were rapid. Much of the fault motion had a strike-slip mode. Alluvial-fan deposits of the New Glasgow and Polly Brook formations (Figure 13) exceed 1 km in thickness adjacent to the Cobequid — Chedabucto Fault Zone (Calder, 1994;

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Figure 13 Stratigraphic framework for the Cumberland Basin (A), with representative pro¢les for the Boss Point Formation (B) from Browne and Plint (1994) and the Joggins Formation (C) from Davies et al. (2005).

Chandler, 1998). The Cumberland Group appears to be absent over wide areas, particularly the Central Carboniferous Platform of New Brunswick, although remnants of the Boss Point Formation suggests that Cumberland strata may originally have been more extensive (New Brunswick Department of Natural Resources, 2000; Johnson, 2003). The onset of this tectonic phase corresponds broadly with the onset of the Alleghanian Orogeny in the U.S. Appalachians, where the Appalachian Foreland Basin was created by a fold-and-thrust belt of Andean elevation (Slingerland and Furlong, 1989; Ettensohn, Chapter 4, this volume). This important period of convergence following collision between Laurussia and Gondwana was manifested predominantly as transtensional and transpressional events in the Maritimes Basin, which lies where the E-W trending Hercynian belt of Europe passes into the NE-SW trending orogenic belt of the U.S. Appalachians (Figure 3; Arthaud and Matte, 1977). Within the Maritimes Basin, rapid subsidence rates, architectural styles, paleoflow patterns, and halokinetic effects collectively testify to the profound influence of tectonism on sedimentation through this period. In the extensional Cumberland Basin (Figure 1), which lies between the fault-bounded Cobequid and Caledonia Highlands, the Mississippian–Pennsylvanian Unconformity is overlain by a predominantly alluvial succession (Figure 13A), much of which is exposed in the remarkable coastal exposures along the Bay of Fundy near Joggins, Nova Scotia. The basal three formations (Boss Point, Little River and Joggins) are about 2,300 m thick and may have been deposited in as little as 2 Myr, suggesting long-term subsidence rates in the order of 1 mm/year (although biostratigraphic boundaries are not clearly defined). The Boss Point Formation, about 800 m thick, comprises alternate braided-fluvial, lacustrine and floodplain intervals, with strong erosion at the base of the fluvial units (Figure 13B). The architecture implies falls and rises of base-level, and abrupt switches of paleoflow suggest tectonic influence (Browne and Plint, 1994). The overlying Little River Formation comprises W600 m of semi-arid to sub-humid redbeds (Calder et al., 2005). The overlying Joggins Formation comprises more than 900 m of stacked cycles that commence with a strong basal transgression marked by an upward transition from coal to limestone to open-water terrigenous deposits with bivalves and ostracodes (Figure 13C). Floral evidence suggests that most of the basin floor was drowned during these transgressions, so that floral elements from the adjacent uplands dominate the open-water facies (Falcon-Lang, 2003a). Subsequent readvance of the coastal plain led to deposition of wetland deposits with spectacular levels of standing trees, followed by alluvial redbeds prior to renewed transgression (Davies and Gibling, 2003; Davies et al., 2005). Joggins architecture is characterized by many flooding surfaces (marked by

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coals, limestones and carbonaceous shales), and the open-water and wetland parts of the cycles are progradational parasequence sets, with thin intervals of retrogradational parasequences below the base of the next cycle. Although the formation contains distributary, meandering and anastomosing fluvial deposits, there is no indication of large valley fills or mature paleosols through this thick stratal interval. This architectural style suggests that base-level rises were enhanced and base-level falls were suppressed as a result of rapid subsidence in the high-accommodation setting of an active extensional basin (Davies and Gibling, 2003). Furthermore, on account of the rapid basinal subsidence, the fluvial channel deposits represent much of the natural geomorphic variability of the original drainage network (Rygel and Gibling, 2006). Coeval strata of the Lancaster Formation crop out at the Fern Ledges near St. John, New Brunswick, where fresh- and brackish-water assemblages include the remains of arthropods, insects and arachnids (Falcon-Lang and Miller, 2007). The overlying red alluvial deposits of the Springhill Mines Formation include thick coals, which onlap the alluvial-fan wedge of the Polly Brook Formation towards the Cobequid Highlands near Springhill (Calder, 1994). The topmost three formations are alluvial redbeds. Fault-bounded basins with coal measures of similar age cover large areas below Prince Edward Island and the Gulf of St. Lawrence and in western Cape Breton (Figure 9; Port Hood Formation: Giles et al., 1997), where they lie along the Hollow Fault Zone and include braided- and meandering-fluvial deposits with a rich assemblage of continental trace fossils (Keighley and Pickerill, 1996, 2003). Coeval alluvial and lacustrine strata are also present in fault-bounded basins in southeastern Cape Breton (Boehner and Prime, 1993) and adjacent to the Cabot Fault system in western Newfoundland (Barachois Group of Bay St. George Basin and Howley Beds of the Deer Lake Basin: Knight, 1983; Hyde et al., 1988), where several formations onlap basement rocks. Probable Barachois equivalents nearly 1,500 m thick were mapped in the St. Anthony Basin (Bell and Howie, 1990).

9.2. Stellarton Group A region of particular interest is the Stellarton Basin, an area of 20 km by 8 km, interpreted as a pull-apart basin at a releasing bend linking the Cobequid and Hollow Faults. Fault motion was mainly dextral along the Avalon — Meguma terrane boundary, located at the site of a right-handed stepover in the Cobequid-Chedabucto fault system (Yeo and Ruixiang, 1987; Naylor et al., 1989; St. Jean et al., 1993; Waldron, 2004). Within the basin area, 3 km of Langsettian to Bolsovian and Asturian strata rest on highly deformed Mississippian rocks, and are lacustrine and fluvial, with coals up to 13 m thick and thick oil shales. Peat (coal precursor) accumulated especially in the northern part of the basin during periods of accelerated subsidence and sediment trapping in the south (Waldron, 2004). The strata show evidence for syn-sedimentary deformation in an extensional mode, and complex fault motions are indicated by patterns of stratal thickening at different levels. Higher members of the formation onlap basement. Thinner, coeval strata north of basin show northward paleoflow, suggesting either that drainage systems bypassed the basin or that the basin was later offset by fault motion. Although the basin fill appears to lack marine fossils, tidal effects are shown by the occurrence of abundant paired mud drapes at some levels (Costain, 2000), and restricted-marine biota are present in the Bolsovian to Asturian Malagash Formation north of Stellarton (Naylor et al., 1998). These observations suggest that tidal influence affected freshwater areas far inland, as in the modern Amazon Basin where tidal effects are experienced at least 800 km inland (Archer, 2005).

9.3. Southwestern maritimes basin In southwestern New Brunswick, Langsettian alluvial fans (Tynemouth Creek Formation) were shed as a result of exhumation at a restraining bend of the Cobequid-Chedaducto Fault (Plint and van de Poll, 1982; Nance, 1987). Within the formation, a remarkable exhumed earthquake fault scarp testifies to syndepositional faulting (Plint, 1985). Evidence of tectonism during this period is also provided by bedrock areas in southwest Nova Scotia (Dallmeyer and Keppie, 1987, 1988; Culshaw and Liesa, 1997; Culshaw and Reynolds, 1997). In this region, a broad zone of deformation overprints the Acadian fold belt, reactivating folds and shear zones. The deformation is dated at about 320 Ma, and accommodated convergence-dominated transpression of the Meguma Zone against an irregular Avalon boundary. Evidence for bedrock thrusting at about 329711 Ma is present near the Cobequid Fault (Waldron et al. 1989).

9.4. Salt migration During this period of active tectonism, the migration of the thick Windsor salt succession, commonly associated with faults, created diapirs in many parts of the Maritimes Basin (Boehner, 1986; Howie, 1988). A large diapir

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Figure 14 Seismic pro¢le 550 through the Cumberland Basin. Note that the Boss Point Formation maintains its thickness over the salt structure at right, whereas the Joggins Formation thins markedly and is present only in the western part of the basin, above the area of a salt weld (W). Modi¢ed fromWaldron and Rygel (2005). Location shown in Figure 1.

field underlies the eastern Gulf of St. Lawrence and the Magdalen Islands (Figure 10), where diapirs are as much as 8 km in height. Diapirs are also prominent in small basins along the Hollow Fault Zone in western Cape Breton, extending across the Cabot Strait to Newfoundland (Langdon and Hall, 1994; Durling et al., 1995). Where seismic lines approach the Cape Breton coast, it is apparent that Windsor Group outcrops lie atop salt structures and that adjacent Mabou and Cumberland group strata form part of drag zones and occupy peripheral basins (Alsop et al., 2000). Windsor diapirs are also prominent on the southern Grand Banks (Figure 12; MacLean and Wade, 1992; Pascucci et al., 1999) and in the St. Anthony Basin. Evaporites are absent from some areas, including much of the Sydney Basin and the Stellarton Basin, where their absence may reflect the local distribution of evaporative basins or salt withdrawal. Salt movement appears to have commenced during the Mississippian (Waldron and Rygel, 2005; Wilson et al., 2006) and to have continued to the latest Carboniferous. Halokinesis predated the formation of the Stellarton Basin where the basal strata are Langsettian (Waldron, 2004), but appears to have terminated prior to the Permian on the Magdalen Islands (Brisebois, 1981). Basal Cretaceous strata on the southern Grand Banks show no influence from underlying Windsor diapirs (Pascucci et al., 1999; Figure 12). Seismic profiles show that the formation-scale architecture of the Cumberland Basin owes much to halokinesis (Figure 14). The Boss Point Formation forms a widespread sheet, whereas the coal-bearing Joggins Formation is present only in the western part of the basin where a salt weld marks the withdrawal of Windsor evaporites that generated nearby diapirs (Waldron and Rygel, 2005). In the Joggins area, salt withdrawal greatly enhanced accommodation creation and promoted high groundwater levels during Joggins deposition, contributing incidentally to the repeated entombment of standing trees in the wetland basins. Lynch and Giles (1996) identified a major decollement (the Ainslie Detachment) near the base of the Windsor Group, and suggested that salt movement at this level had accommodated regional extension, as well as accounting for some apparently disjunct stratigraphic relationships. However, many unusual stratal relationships at the Windsor Group level may reflect the formation of individual salt structures and welds (Alsop et al., 2000; Waldron and Rygel, 2005), or local gravity sliding during Windsor deposition (Thomas et al., 2002).

10. Pennsylvanian-Permian: Thermally Subsiding Basins and Second Phase of Coal Measures 10.1. Thermal subsidence As the Carboniferous proceeded, most parts of the Maritimes Basin subsided more slowly, and alluvium of the Upper Cumberland Group and the Pictou Group onlapped basement rocks over wide areas, particularly across the Sydney and Gulf of St. Lawrence basins, over bedrock at the basin margins, and across the Central Carboniferous Platform of New Brunswick (Figure 1). The broadly elliptical shape of the modern Maritimes Basin (Figure 1), centered on the Gulf of St. Lawrence, largely reflects the distribution of strata within this thermally subsiding region, linked in turn to the basin’s position within the Que´bec Reentrant and between promontories of the former continental margin. This change in basinal style reflects the increasing importance of thermally driven subsidence across the region, and suggests waning tectonic activity, although up to 4 km of strata accumulated locally during this period (Figure 9).

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Figure 15 Paleogeographic map of part of eastern Canada during the Pennsylvanian. Flow directions based on paleo£ow data summarized in Gibling et al. (1992); Ages for paleo£ow data: L, Langsettian; B, Bolsovian; S, Stephanian; P, Permian. Areas of some major coal¢elds are shown, ranging in age from Langsettian to Asturian. Modi¢ed from Atlantic Geoscience Society (2001).

Regional paleoflow data (Figure 15) suggest that the basin was traversed by a large river system with headwaters in the rising Appalachian mountain chain to the southwest, draining eastward into the MidEuramerican Sea (Gibling et al., 1992). The widespread alternation between sandstone- and shale-rich formations manifested in the Gulf of St. Lawrence succession suggests a link to large-scale tectonic activity, probably in the headwaters of the mountain belt (Rehill, 1996). Correlation of individual coals (Haites, 1952; Hacquebard, 1986) and formations (see below) between the Sydney and Gulf successions suggests that, late in the Carboniferous, large areas of the Maritimes Basin had coalesced into a single basin, partially or completely covering former uplands, and some individual coals may have covered much of this area, which was similar to that of the U.S. Illinois Basin (Gibling et al., 2004). Ryan and Boehner (1994) suggested that the Cumberland Basin succession was originally 3–4 km thicker than presently observed. Similar strata may have covered most bedrock in Atlantic Canada, prior to a major period of exhumation and erosional stripping commencing in the latest Paleozoic, probably after about 280 Ma (Hendriks et al., 1993; Ryan and Zentilli, 1993; Ryan and Boehner, 1994).

10.2. Morien and Pictou groups In the Sydney Basin, much of the Pennsylvanian is represented by the 2,500 m thick Morien Group. Sedimentation did not recommence until the Duckmantian, after which strata of the South Bar Formation covered the Mississippian–Pennsylvanian Unconformity (Figures 7 and 8). These braided-fluvial sandstones occupied broad valleys on the unconformity surface and include rare thin coals and reworked peat mats with a few prominent sequence boundaries (Tibert and Gibling, 1999; Figure 16). In contrast, the overlying Sydney Mines Formation is shale- and coal-rich, and comprises stacked sequences and sequence sets with a predictable facies organization and a mean duration in the Milankovitch band (Figure 16; Gibling and Bird, 1994; Batson and Gibling, 2002; Gibling et al., 2004). Lowstand surfaces are marked by valley fills, calcretes and vertisols (Tandon and Gibling, 1997), and maximum flooding surfaces are marked by coals and bivalve-ostracode limestones, with some valleys probably cut and filled during falling stage. The high sulfur content of the coals and the presence of

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Figure 16 Model for the coal window in the cratonic Sydney Basin, with representative stratigraphic architecture for the South Bar and Sydney Mines formations. Modi¢ed from concepts in Calder and Gibling (1994) and Gibling et al. (2004). SB ¼ sequence boundary; TSTand HST ¼ transgressive and highstand systems tracts, respectively; mfs ¼ maximum £ooding surface.

restricted-marine, agglutinated foraminifera (Gibling and Wightman, 1994) and glaucony indicates a marine connection for the basin. The presence of calcrete on lowstand surfaces implies relative aridity during periods of lowered sea-level, as in many regions during the Last Glacial Maximum of the Quaternary. Within Sydney sequences, Falcon-Lang (2004) documented a contrast between cordaitalean-dominated floras of valley fills and dryland alluvial plains, which accumulated during glacial lowstand and early transgressive periods, and lycopsid-dominated floras of interglacial late transgressive and highstand periods, during which tropical rainforests expanded from their dry-phase refugia during the cool, dry glacial periods. Based on available biostratigraphic information, sedimentation rates for much of the succession were less than 0.2 mm/year (Gibling et al., 2004) — less than 20% of the estimated subsidence rate for the Cumberland Basin fill. Within this low-accommodation setting, glacioeustatic fluctuations generated by Late Paleozoic glaciation would have exerted a strong influence on stratal architecture (Gibling and Rygel, in press), and peat accumulation

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largely reflects high-accommodation events generated during short periods of rapid sea-level rise (Gibling et al., 2004). Capping redbeds of the Pictou Group comprise the topmost Carboniferous and probably the Permian. The distribution of coal within the basin fill (the ‘‘coal window’’ of Calder and Gibling, 1994; Figure 16) reflects a basin hydrology suitable for peat accumulation, high-accommodation events that promoted peat accumulation and preservation, and a progression to more arid climates that eventually precluded peat accumulation. Below Prince Edward Island and the Gulf of St. Lawrence, the younger Carboniferous succession is known largely from well and seismic information, and age dates are based mainly on palynological analysis. Rehill (1996) and Giles and Utting (1999) identified four formations (currently informal) with B4 km of aggregate thickness above the Early Westphalian Unconformity, resting on the Port Hood Formation of the Lower Cumberland Group in the east and extending across Mississippian strata onto basement rocks to the west (Figure 9). The basal Bradelle formation thickens towards Cape Breton Island and is dominated by sandstone, with gray and minor red shales and numerous coals and coaly shales. The overlying Green Gables formation is shale- and coal-rich, whereas the Cable Head and Naufrage formations comprise sandstone, red shale and pedogenic carbonate. These strata are broadly equivalent to formations in the Cumberland Basin succession in northern Nova Scotia (Figure 17). The lower three formations of the Gulf succession and the coeval Malagash and Inverness formations (Bolsovian to Asturian) along the southern Gulf margin have yielded agglutinated foraminifera that provide evidence for restricted-marine influence (Wightman et al., 1994; Brown, 1998; Naylor et al., 1998). In comparing the Sydney and Gulf of St. Lawrence successions, age dates and lithology suggest correlation between the fluvial sandstones of the Bradelle and South Bar formations, the coal-rich Green Gables and Sydney Mines formations, and the predominantly red Cable Head and Naufrage formations and the Pictou Group (Figure 17). Although no definitive Permian dates have been obtained, the uppermost exposed strata in Nova Scotia and on Prince Edward Island probably extend into the Permian (van de Poll, 1989; Calder, 1998). A good tetrapod trackway record is present at several localities (Mossman and Place, 1989; Van Allen et al., 2005; Calder et al., 2004). The Gulf of St. Lawrence succession extends westwards across the Central Carboniferous Platform and Moncton Basin of New Brunswick, where the Pictou Group is about 1 km thick and rests on a regional disconformity or unconformity above the Mabou Group, which it oversteps onto older bedrock (New Brunswick

Figure 17 Summary correlation chart for Duckmantian to Permian strata of the Maritimes Basin. See Figure 1 for locations. Dashed lines on columns indicate probable age of formation boundaries where reasonably well established; in other columns, formation names are placed opposite the most probable age. Palynomorph dates for basal strata in most areas lie within a Late Duckmantian/Early Bolsovian range. Topmost strata in the Western Maritimes Basin and Magdalen Islands are undated but probably Permian. Sources of information: Bathurst area: Ball et al. (1981), van de Poll (1995); Minto area: St. Peter (1993), van de Poll (1995), Kalkreuth et al. (2000); Moncton Basin: Johnson (1995), St. Peter (2001); Cumberland Basin: Ryan and Boehner (1994); Western Maritimes Basin: Rehill (1996), Giles and Utting (1999); Magdalen Islands: Brisebois (1981),Tanczyk (1988); Western Cape Breton Island: Giles et al. (1997); Sydney Basin: Pascucci et al. (2000). Radiometric ages from Hess and Lippolt (1986) and Menning et al. (2000).

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Department of Natural Resources, 2000). Across this region, the Pictou Group is considered to include all strata above a Late Duckmantian to Early Bolsovian unconformity (C. St.Peter, personal communication, 2005), and includes the important Minto Coal of Bolsovian age (Kalkreuth et al., 2000). Palynological analysis of these predominantly redbeds is difficult, but breaks may be present within the Pictou Group at some levels (C. St. Peter, personal communication, 2005). In seismic profiles, coal-measure intervals gradually lose their distinctive character westward from the Gulf towards New Brunswick, where they appear to pass into terrestrial strata without coal. In western Newfoundland, Bolsovian strata rest on basement, and alluvial-fan deposits have yielded spectacular examples of reworked upland vegetation in the form of petrified cordaitalean logs nearly 2 m in diameter – the remains of early conifers that were almost 50 m high (Falcon-Lang and Bashforth, 2004, 2005).

11. Pennsylvanian to Permian Sedimentation and Tectonic Events: Late Stages of Pangean Assembly 11.1. Permian eolian sandstones Below Prince Edward Island and the eastern Gulf of St. Lawrence, 700 m of sandstone with minor red shale were termed ‘‘Unnamed Permian Sandstones’’ by Giles and Utting (1999), and appear to overlie a profound local unconformity (Figure 9). On the nearby Magdalen Islands, the Cap-aux-Meules Formation includes nearly 700 m of eolian sandstones of probable Permian age, resting on Windsor strata in outcrop and on the Pictou Group in offshore wells (Brisebois, 1981). The two rock units probably represent a distinct Permian phase of accumulation under arid conditions. The Gulf of St. Lawrence sandstones were penetrated in a series of wells drilled on or close to salt structures (Howie, 1988, Figure 10), and the Magdalen Island sandstones overlie diapirs. The locally deeply erosive base of the unit may reflect exhumation linked to salt tectonics.

11.2. Final tectonic events Convergence associated with Pangean assembly continued in the central and southern U.S. Appalachians, as well as in some parts of Atlantic Canada, during the latest Carboniferous and Permian. In the Sydney Basin, modest differential subsidence is recorded over basement structures (Gibling et al., 2002). More significantly, a phase of transpressive deformation reactivated tectonic lineaments and affected the Morien and Pictou groups (Figure 8), probably reflecting a little-known Permian phase of deformation (the Donkin Episode: Pascucci et al., 2000; Gibling et al., 2002). The Stellarton Basin experienced late-stage shortening following extensional events, including generation of a positive flower structure, possibly in the Permian (Waldron, 2004), and Jutras et al. (2003) documented Alleghanian faulting, possibly of late stage, in Gaspe´. This activity may have been linked to rotation of lithospheric blocks during the final stages of Pangean assembly (Faure et al., 1996; Vai, 2003). Such late-stage deformation, as well as later Mesozoic structural events, may have influenced hydrocarbon trapping in some areas, for the majority of organic maturation took place much earlier, during maximum burial in the Pennsylvanian and earliest Permian, and later deformation may have breached some reservoirs (Pascucci et al., 2000; Gibling et al., 2002). Srivastava and Verhoef (1992) used geophysical data to reconstruct the final configuration of eastern Canada and western Europe prior to Atlantic rifting. Their maps place the Grand Banks adjacent to northwest Africa and Iberia. Although paleogeographic reconstructions for this region have been attempted (Ziegler, 1989), the correlation of tectonic lineaments, terranes and stratal groups between Atlantic Canada and adjacent areas is problematic, and no comprehensive model currently explains the links between the two areas. Correlations are impeded because crucial areas of junction lie below the modern, offshore continental margin, and perhaps because Mesozoic to Cenozoic Atlantic opening reactivated fundamental tectonic lineaments between the Grand Banks and Iberia, obscuring their earlier history. Lefort et al. (1993) suggested that southwest-dipping features imaged in deep-crustal seismic lines across the Grand Banks are a continuation of the Hercynian nappes of southwestern Europe.

11.3. Climate change Calder (1998, Figure 5) presented a climate curve for the Maritimes Basin fill. As indicated by the presence of calcretes, vertisols and evaporites, semi-arid conditions prevailed during deposition of the Mississippian Horton, Visean and Mabou groups, continuing into the basal Namurian to Langsettian strata of the Cumberland Group. Subsequent more humid conditions are indicated by hydromorphic paleosols and coals, although climates

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probably remained strongly seasonal. Calcretes and vertisols, with eolian beds in the youngest basin fill, mark a return to drier conditions in the Asturian through to the Permian. The Maritimes Basin curve corresponds with a change in the U.S. Appalachian Basin from relatively arid Mississippian conditions to more humid Pennsylvanian conditions with peat accumulation, to more arid topmost Carboniferous and Permian conditions (Phillips et al., 1985; Witzke, 1990). Such regional climatic change probably represents northward plate drift from the Late Devonian through to the Late Permian, which progressively shifted eastern North America from the southern hemisphere dryland zone in the Mississippian to the equatorial zone in the Pennsylvanian, to the northern hemisphere dryland zone by the later Permian (Schutter and Heckel, 1985; Torsvik and Cocks, 2004). Additionally, monsoonal conditions would have developed as Pangea assembled (Kutzbach and Gallimore, 1989).

12. Permian to Mesozoic: End of Maritimes Basin Deposition, and the Breakup of Pangea Pangea was not stable for long. Rift basins developed along pre-existing faults such as the Cobequid Fault, and were filled with thick Triassic and Jurassic sedimentary and volcanic successions. Although no age dates precisely define the onset of rifting in Atlantic Canada, Olsen et al. (2000) suggested a Permian commencement, based on correlation between stratal successions in rift basins in New Brunswick and Morocco. Rifting was accompanied by regional exhumation, as indicated by fission-track evidence for Triassic cooling (Gibling et al., 2002), and erosion of considerable thicknesses of Carboniferous to Permian strata. A more widespread period of sedimentation began during the Cretaceous as rifting progressed to seafloor spreading, with local deposits laid down across Carboniferous rocks onshore and offshore in many parts of Atlantic Canada (Figure 12). In parts of mainland Nova Scotia, Cretaceous rivers flowed across a karstic terrain developed by prolonged weathering of Windsor group carbonates and evaporites (Falcon-Lang et al., 2007).

13. Synopsis of Maritimes Basin History 13.1. Key points Maritimes Basin history shows several key features: 1. The collision of terranes and larger continental elements (Laurussia, Gondwana) was largely completed during the Acadian Orogeny, prior to or during the earliest part of the basin’s history. Subsidence and tectonic activity continued for B120 Myr thereafter. 2. Continental-scale fault zones, some with strike-slip mode, were active through much of the basin’s history, and controlled sedimentation locally. 3. Numerous phases of widespread or more local basin subsidence and inversion took place. Rapid alternation of subsidence and inversion implies a series of extensional and compressional periods, mostly affecting smaller depocenters. Extension and compression appear to have been periodically coincident in local depocenters in different parts of the region, especially during the Mid-Devonian to Tournaisian. The duration of subsidence phases was a few million years to a few tens of millions of years. Such a dynamic environment would have caused both creation and destruction of hydrocarbon reservoirs, if conditions for maturation were suitable. 4. The most active phases of basin subsidence represent extension along bounding faults, enhanced locally by salt withdrawal. In addition, much of the region was affected by two periods of thermally driven subsidence, during which fault activity was less prominent. The two types of subsidence yielded distinctive styles of stratigraphic architecture, and the basins filled largely with continental strata of alluvial-fan, alluvial plain, lacustrine and coastal facies. 5. Marine sediments are present at some levels, but seaways were restricted, as indicated by thick evaporites, the scarcity of open-marine biota, and the predominance of restricted-marine biota. Coastal and marine sequences at several levels show evidence for Milankovitch control. 6. Volcanic activity (bimodal) and plutonic activity (largely granitic) were concentrated in the earlier period of basin history but continued episodically. 7. Long-term climatic variations were linked primarily to the region’s sustained northward drift. At the largest scale, Atlantic Canada lay within a zone of dextral mega-shear during the terminal collision of Gondwana and Laurussia (Arthaud and Matte, 1977). Oblique convergence accounts for the strike-slip fault

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system in the environs of the Maritimes Basin (Figures 12 and 13 in Chapter 1), passing southwestward into an extensive region of crustal thickening associated with Pennsylvanian to Permian deformation of the eastern U.S. Appalachians. The transition to this zone of crustal thickening may be encountered in the reworked zones of southwest Nova Scotia, which were being ‘‘transpressed’’ in the Carboniferous at ca. 320 Ma (Culshaw and Reynolds, 1997). The deformation style was ductile, and implies greater crustal thickness than in areas associated with basin development. Maritimes Basin history occupied as much as B120 Myr during convergence that followed terrane accretion and ocean closure in the Acadian Orogeny, and most of the basin filling (Horton to Pictou groups, Figure 4) took place over 60–80 Myr, from the Late Devonian to the Early Permian. The process of closure and collision was variable along strike and probably strongly diachronous. Through this period, the orogen lay within a continuously evolving dynamic boundary within a newly formed continental plate, and long-continued convergence must have affected both crustal and mantle lithosphere. One implication of this setting is that uppercrustal processes such as basin formation and faulting should have been linked with magmatism and thermal effects originating in the converging lithospheric and asthenospheric mantle. We postulate that the extended history of the Maritimes Basin, the style of individual basins, their relationship with strike-slip faulting, periodic inversion, and association with mafic and felsic magmatism find their origin in the dynamic environment of postcollisional convergence, in which the fate of the crust and mantle lithosphere are intimately linked. Thus, a lithosphere-scale approach (Keppie and Dallmeyer, 1995) may be useful for understanding Late Paleozoic events in Atlantic Canada, and would account for sedimentation, magmatism, and deformation under a single umbrella. Aspects of the late development of the Variscides of Europe have been explained in a similar way (Matte, 1991). Along tectonic strike from the Maritimes Basin in the zone of convergence, they have a CarboniferousPermian history of basin formation and similar bimodal magmatism. Henk (1999) showed that the diverse sedimentary and magmatic phenomena could be explained as a result of lithospheric extension caused by a combination of gravitational forces and tectonic stresses — resulting, for example, from strike-slip faulting. Meissner (1999) interpreted reflection seismic data from the Variscides as evidence for extension in the lower crust, implicitly supporting Henk’s hypothesis.

13.2. Magmatic history How have igneous events in the history of the Maritimes Basin been explained, and how do they accord with this scenario of prolonged convergence? Major igneous phases have been documented in the Mid-Devonian, Late Devonian, latest Devonian to Early Carboniferous, and Late Tournaisian to Early Visean, with minor Pennsylvanian activity (Barr et al., 1985; Pe-Piper and Piper, 1998, 2002; Dessureau et al., 2000). Devonian volcanics are present in New England (Thompson and Hermes, 2003). The mafic volcanics are predominantly continental tholeiite flood basalts, with a few younger alkalic volcanics. Stacked basalt flows are up to 1.5 km thick (Dessureau et al., 2000), and the volcanic suite is bimodal, with rhyolites prominent in many volcanic successions. Possible explanations adduced for these volcanic periods include mantle plumes, and rapid lithospheric thinning associated with crustal extension, during which magmas may have been emplaced along faults (Durling and Marillier, 1990; Dessureau et al., 2000; Pe-Piper and Piper, 2002). Marillier and Verhoef (1989) used geophysical evidence to infer that an underplated layer 10–20 km thick underlies part of the Maritimes Basin. The South Mountain Batholith may have formed in a Late Devonian continental-margin arc to syncollisional setting during the docking of the Meguma Terrane (Tate and Clarke, 1995; Clarke et al., 1997) or during subsequent convergence (Murphy et al., 1999). The main central plutons have been taken to imply crustal thickening (Keppie and Dallmeyer, 1995), whereas the peripheral plutons with more associated mafic material (Tate and Clarke, 1995) may suggest crustal melting and hybridization, with mafic melts supplying heat (Clarke et al., 1997). Murphy et al. (1999) suggested that Acadian and later convergence overrode a mantle plume, analogous to Laramide events in western North America. A more comprehensive view of Maritimes Basin igneous activity draws upon analogues in orogenic belts of varied age. The common occurrence throughout geological time of bimodal magmatism in regions of postcollisional convergence, as well as late tectonic ‘‘granite blooms’’, seems to require a common mechanism that finds its origin in the ‘‘normal’’ behavior of continental lithosphere during and after collision, rather than requiring episodic plume activity or being restricted largely to extensional zones associated with faults. In zones of post-collisional convergence, the lithosphere commonly experiences delamination, in which surficial crustal material is separated from underlying mantle lithosphere. Sacks and Secor (1990) and Nelson (1992) inferred that delamination played an important role in the U.S. Appalachian Orogen, and Keppie and Dallmeyer (1995) suggested that delamination contributed to Maritimes Basin history. Delamination potentially allows large volumes of mafic magma to be generated, with widespread heating of the overlying crust. During the Devonian

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and earliest Mississippian part of the basin’s history, delamination could have resulted in considerable volcanic and plutonic activity without requiring excessive crustal thickening. Extensional effects, including upper-crustal faulting and basin formation, would also have accompanied delamination (Sacks and Secor, 1990). Additionally, Nelson (1992) noted that reestablishment of the lithosphere subsequent to delamination and orogenic collapse produces a thermal ‘‘sag’’ basin on top of the collapsed orogen. The two near-basinwide periods of thermal subsidence recorded in the Maritimes Basin may reflect the regional response of the lithosphere to major periods of convergence, delamination, and orogenic collapse. Magmatic and thermal events continued in the region through to the Permian, especially in southwest Nova Scotia (Pe-Piper and Loncarevic, 1989) which lay close to well-documented Permian magmatic activity in New England (Wintsch and Sutter, 1986). A 252 Ma date for the Malpeque Dyke of Prince Edward Island (Greenough and Fryer, 1991) may be linked to the early stages of Pangean rifting, which was accompanied by volcanism and possible plume activity in Europe and northwest Africa (Doblas et al., 1998).

14. A Modern Analogue: Turkey and Eastern Mediterranean The Late Cenozoic tectonic development of Turkey and environs furnishes a modern analogue to the Late Paleozoic Maritimes Basin system (Figure 18). Points of similarity include their position within a large-scale orogenic system, post-collisional (syn-convergence) magmatism, basin formation, and strike-slip faulting (S- engo¨r, 1979; S- engo¨r and Yilmaz, 1981; S- engo¨r et al., 1985). Additionally, both change along tectonic strike from a region with strike-slip faults to a fold-and-thrust zone — in the case of Turkey, from the Anatolian region to the Zagros region of Iran. Basin formation in Turkey is evidently part of a whole-lithosphere system, given the combination of upper-crustal faulting, crustal thinning, sedimentation coeval with tectonic activity, and mafic magmatism. The comparison strengthens arguments for a lithosphere-scale understanding of the Maritimes Basin. One of the more obvious shortcomings of the Turkish analogue is that the present tectonic regime began at ca. 20 Ma and is ‘‘in progress’’. Thus, its history has been of much shorter duration than the 120 Myr history of the Maritimes Basin, and the Turkish analogue may correspond best with the earlier part of Maritimes history. The Early Miocene in Turkey includes closure of two sutures within Anatolia (Inner Tauride and Izmir-AnkaraErzincan, S- engo¨r et al., 1985). Crustal thickening occurred throughout Anatolia during this stage, which may be compared to Acadian (ca. 390 Ma) terrane accretion, and took place in both orogens before the main period of basin development. Mid- to early Late Miocene closure of the Bitlis suture (Figure 18) resulted from the collision of Arabia with Eurasia, and may be analogous in its effects to the collision of Gondwana with Laurussia, although smaller in scale. Marine incursions persisted locally in Anatolia, and the Aegean began to develop. Continuing post-collisional convergence of Arabia accounts for ongoing crustal thickening in eastern Anatolia (Eastern Anatolia Collision Zone (EACZ), Figure 18) and induced ‘‘tectonic escape’’ of Central and Western Anatolia toward the eastern Mediterranean subduction zone (Dewey et al., 1986). Crust thins westward from eastern Anatolia (50 km) to reach a minimum in the Aegean (S- engo¨r et al., 1985). The Aegean is underlain by thinned continental crust and may properly be thought of as a submerged continuation of Western Anatolia. The Northern and Eastern Anatolian faults (NAF and EAF, Figure 18), the boundaries of the expelled Anatolian block, are strike-slip faults that exerted important controls on the terrestrial sedimentation that began soon after closure across the Bitlis suture. Basins within the interior of the Anatolian block are poorly exposed in central Anatolia (OVAS terrain, Figure 18), where they form fault-bounded depressions with Neogene strata and volcanics (S- engo¨r, 1979), but basins are prominent in western Anatolia where they are confined within welldeveloped graben formed in the ‘‘stiff lid’’ of the actively extending and thinning lower crust. Although fewer in number than in the west, there are also basins east of the expelled Anatolian block within the currently thickening eastern Anatolian crust (Figure 18). Magmatism following terrane accretion in Western Anatolia at ca. 20 Ma is attributed to lithospheric mantle melting due to delamination (Aldanmaz et al., 2000; Aldanmaz, 2002). This magmatism may be compared to Late Devonian magmatism in the Meguma Terrane (South Mountain Batholith and associated intrusions). Magmatism has continued from ca. 11 Ma until the Holocene in Western Anatolia (Aldanmaz et al., 2000) and since ca. 6 Ma in Eastern Anatolia (Pearce et al., 1990). In both areas, magmatism has been attributed to melting of the lithospheric mantle, linked to lithospheric extension in the west and delamination accompanying convergence in the east. This later magmatism, accompanying basin formation, may be compared with Late Devonian to Early Carboniferous magmatism in the Maritimes Basin. From a sedimentary viewpoint, there are also many points of similarity between modern Turkey and the Maritimes Basin. Lake deposits are prominent in Turkish strike-slip basins (Dunne and Hempton, 1984), comparable to parts of the Horton Group, and tectonically active regions are accumulating fluvial and deltaic sediments (Aksu et al., 1987). Convergence has resulted in the progressive loss of marine connection and the

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Figure 18 (a) Simpli¢ed geology of Turkey and adjacent areas. AEZ, Aegean extensional zone; EACZ, East Anatolian Collisional Zone; DSF, Dead Sea Fault; EAF, East Anatolian Fault; NAF, North Anatolian Fault; OVAS, region of fault-bounded depressions with Neogene strata and volcanics (S-engoºr, 1979); AEZ,West Anatolian Extensional Zone. (b) Maritimes Basin and adjacent areas at similar scale to a; see also Figures 1 and 3.

establishment of restricted-marine embayments such as the Black Sea (Figure 18; Degens and Ross, 1974), where deep, anoxic waters are capped by low-salinity water derived from the Danube River. The Mediterranean is also a closing ocean, and its Late Cenozoic sediments include thick evaporites and salt structures, suggesting broad comparison with the Windsor Group. Although the region is not noted for its peat accumulations, peats are associated with lacustrine deposits in the Hula Valley on the Dead Sea Fault in Israel (Picard, 1965), perhaps comparable to Stellarton Basin coals. From the preceding account, it can be seen that, at the largest scale, the Zagros-Anatolia area resembles the Alleghanian Orogen–Maritimes Basin region of the eastern United States and the Atlantic Provinces. The correspondences include: (1) crustal thickening in the Zagros, comparable to Alleghanian deformation in eastern United States, (2) crustal thickening in Eastern Anatolia (at the transition from thickening to thinning crust), comparable to ductile Carboniferous deformation in southwest Nova Scotia, (3) fault-bounded basins of Western Anatolia, comparable to depocenters within the orogenic zone of the Maritimes Basin, (4) Northern and Eastern Anatolian and other basin-bounding faults, comparable to strike-slip faults of the Maritimes Basin, and (5) magmatism originating in the asthenospheric or lithospheric mantle in both regions. Periods of regional thermal subsidence are not apparent in the Late Cenozoic history of Turkey and environs, but have been documented in the Pannonian and Transylvanian basins of eastern Europe. In the Pannonian Basin, the Great Hungarian Plain has experienced post-rift subsidence over the past 10 Myr (Huismans et al., 2001, Figure 5D). In the Early Miocene of the Transylvanian Basin, the structural style of basin evolution changed from

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one of active, fault-related subsidence to a regime of strong subsidence that took place with only local activity of major faults, allowing about 5 km of strata to accumulate in a 15–20 million year period (Huismans and Bertotti, 2002). Although these basins lie in a back-arc setting, they may provide a partial analogue for the two regional periods of thermal subsidence experienced by the Maritimes Basin, each persisting through a few tens of millions of years. Additionally, different parts of these eastern European basins have experienced simultaneous extension and contraction.

ACKNOWLEDGMENTS This paper has benefited enormously from comments and discussion with many colleagues over a prolonged period, especially Sandra Barr, John Calder, Peter Giles, Rob Naylor, and Clint St. Peter. N. Culshaw thanks David Piper for discussions about the geology of Turkey and environs. We are grateful to Ashton Embry and Andrew Miall for their comments on an earlier version of the manuscript, and to Andrew Miall for editorial assistance. We thank Sandra Barr for providing Figure 2, and Peter Giles for providing an early version of Figure 9. Funding was provided from Natural Sciences and Engineering Research Council of Canada grants to M. Gibling and N. Culshaw, and formative ideas arose during V. Pascucci’s tenure of a Killam Post-doctoral Fellowship at Dalhousie University. Recent research at Joggins, referred to in the text, was supported by grants from the American Chemical Society (Petroleum Research Fund) and Imperial Oil.

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CHAPTER 7

Pennsylvanian–Jurassic Sedimentary Basins of the Colorado Plateau and Southern Rocky Mountains Ronald C. Blakey

Contents 1. Introduction 1.1. Location 1.2. Stratigraphic interval 1.3. Scope and organization 2. Precambrian Basement and its Possible Control on Phanerozoic Deposition 2.1. Trends and lineaments 2.2. Younger Precambrian sedimentary basins 3. Phanerozoic Tectonics and Depositional History 3.1. Early and Middle Paleozoic 3.2. Pennsylvanian–Permian 3.3. Triassic 3.4. Jurassic 3.5. Cretaceous 3.6. Cenozoic 4. Pennsylvanian–Middle Jurassic Sequence Stratigraphy 4.1. Introduction 4.2. Pennsylvanian 4.3. Permian 4.4. Triassic 4.5. Jurassic 5. Tectonic Origins of Pennsylvanian–Permian Basins 5.1. Introduction 5.2. Yoked basins 5.3. Non-yoked basins 5.4. Cordilleran basins 6. Tectonic Setting of Triassic Basins 6.1. Introduction 6.2. Moenkopi shelf 6.3. Eastern Cordilleran Basin 6.4. Pre-Shinarump paleovalleys and Shinarump deposits 6.5. Chinle Basin 7. Tectonic Setting of Jurassic Basins 7.1. Introduction 7.2. Zuni sag 7.3. Utah–Idaho trough 8. Summary: Tectonic Evolution and Controls on Deposition 8.1. Tectonic sequence of events 8.2. Climatic controls 8.3. Eustatic controls 8.4. What was the dominant control? Acknowledgments References

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Abstract Pennsylvanian through Jurassic sedimentary basins of the southern Western Interior of the United States are present within the Colorado Plateau and Southern Rocky Mountains provinces. Until the latest Mesozoic to Early Cenozoic, the two provinces were mostly indistinguishable; from latest Precambrian through Cretaceous they formed the SW margin of the North American craton and both were at or very near sea level. From the Pennsylvanian through the Jurassic, a series of sedimentary basins formed throughout the region. Unlike many other cratonic basins of North America, the basins covered in this chapter were each relatively short-lived as the location, shape, style, and nature of sedimentary fill constantly evolved. Late Paleozoic basins were related to the complex tectonic evolution of the Ancestral Rocky Mountains; basins adjacent to basement-cored uplifts preserve thick clastic wedges of coarse arkosic sedimentary rocks that grade distally into marine mudstone and carbonate. Basins more distal to uplifts were filled with mixtures of shallow marine and continental deposits. Both proximal and distal basins commonly have evaporite deposits that reflect warm, arid, tropical conditions. Triassic and Jurassic basins bear little resemblance to those of the Pennsylvanian and Permian. Mesozoic basins of the southern Western Interior were related to complex tectonic events of the evolving Cordilleran arc and coeval Mesozoic terrane accretion to western North America. The thickest basins are elongate N-S along the western margin of the Colorado Plateau but are abruptly truncated westward by pre-Cretaceous erosion. Lower Triassic rocks are marine to the west and continental fluvial to the east. Upper Triassic rocks are exclusively continental. Jurassic rocks are dominantly continental and contain some of the most extensive eolian deposits in the stratigraphic record; Middle and Upper Jurassic rocks grade westward into marginal marine deposits.

1. Introduction 1.1. Location From the Late Precambrian through the Jurassic, the Colorado Plateau and Southern Rocky Mountains regions lay on the southwest margin of cratonic North America. Although now distinct physiographic provinces, the two regions had similar geologic histories throughout much of this span of time. This chapter covers the Colorado Plateau and vicinity and Southern Rocky Mountains Region including all of the states of Arizona and Utah, eastern Nevada, southern Wyoming, and most of New Mexico and Colorado, except their easternmost portions (Figure 1). The region is everywhere underlain by Precambrian crystalline basement that was welded onto North America by 1.4 Ga. The veneer of Paleozoic and Mesozoic sedimentary rocks, the subject of this chapter, mantles most of the region except where stripped by erosion across Laramide Cenozoic uplifts. These strata were deposited on wide shelves and in sedimentary basins that flanked cratonic North America and lay east of the paleo-Pacific Ocean. Paleozoic and Lower Mesozoic sedimentary rocks were deposited during three distinct tectonic episodes that mark the transition of western North America from a passive margin to an active margin. Early and Middle Paleozoic strata formed on a stable passive margin that succeeded the rifting of North America from Gondwanaland during the Late Precambrian. Mississippian, Pennsylvanian, and Permian rocks were deposited during a time of transition and significant regional orogenic activity, and Triassic and Jurassic rocks formed in a general back-arc setting behind the rapidly building and evolving Cordilleran Arc (Figure 2). The sedimentary rocks east of the Cordilleran (Wasatch) hingeline have experienced minor to major post-depositional orogenic disruption but all lie at or very close to their original site of deposition and no palinspastic restoration is necessary. Rocks west of the hingeline require extensive palinspastic restoration for both Mesozoic shortening and Cenozoic extension. The sedimentary rocks of the region are exposed in a variety of high desert and intermontane settings. On the Colorado Plateau, exposures vary but are dominated by continuous, well-exposed, undeformed outcrops on which three-dimensional studies can be performed. The Southern Rocky Mountains consists of poor to excellent outcrops that display some degree of tectonic deformation and locally discontinuous exposure. Both regions have been regionally uplifted over 2,000 m during the Tertiary. The Colorado River and its tributaries have carved the landscape into its famous scenic, dissected, landforms during the last 5–10 million years. The extreme local relief, commonly approaching 2,000 m, and local structure expose the complete section or major portions of the Paleozoic through Tertiary section across fairly small areas.

1.2. Stratigraphic interval The emphasis of this chapter is the stratigraphic record from the Pennsylvanian through the Middle Jurassic. Late Precambrian and Early and Middle Paleozoic rocks are briefly covered; Cretaceous and Tertiary rocks of the region are covered in a separate chapter to this volume. In general, the Phanerozoic section thickens irregularly

Pennsylvanian–Jurassic Sedimentary Basins of the Colorado Plateau

Figure 1 text.

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Map of Colorado Plateau and Southern Rocky Mountains and adjacent regions showing features mentioned in the

westward to over 3,000 m across the central and western Colorado Plateau. A wide variety of shallow marine, shoreline, and continental deposits are present, but the region is most famous for its Permian–Jurassic fluvial and eolian strata. Regional correlation is based on published literature and my own stratigraphic work and, in general, tends to conform to nomenclature in use by the U. S. Geological Survey, although a few correlation problems persist, mostly at a detailed level beyond the scope of this chapter.

1.3. Scope and organization The chapter is organized as follows: (1) Summary of tectonic and depositional history sections relate Precambrian and Phanerozoic tectonic settings to controls on Phanerozoic deposition; (2) these are followed by presentation of Pennsylvanian through Middle Jurassic sequence stratigraphy with emphasis on sedimentary accumulations and

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Figure 2 Paleotectonic settings of southwestern North America. A, Late Precambrian to Mississippian; B, Pennsylvanian-Permian; C,Triassic; D, Jurassic. Heavy solid line, western margin of North America; heavy line with ticks, convergent plate boundary-subduction zone, ticks on upper plate; heavy dashed lines, hypothetical rifted blocks of North America; light blue, thin deposits; medium blue, thick deposits, basins; orange with dots, Late Precambrian marginal basins; deep blue, ocean crust; pink, exposed Precambrian crust; green, arcs, terranes; red, Cordilleran arc; yellow, Mesozoic eolian ergs; WL,Wasatch line; MoP, Mojave province; YP,Yavapai province; MaP, Mazatzal province; JL, Jemez lineament; GP, Grenville province; SA, Sonoman arc; RMT, Roberts Mountain thrust; ARB, Ancestral Rockies basins; ARU, Ancestral Rockies uplifts; Hav B, Havallah basin; M--CC, McCloud arc--Cache Creek fore arc; MA, McCloud arc; Utr BA, Upper Triassic back arc basin; GT, Golconda thrust; LTr, Lower Triassic; M, marine; Utr, Upper Triassic; RAR, remnant Ancestral Rockies; Mog Sl, Mogollon slope; LJBA, Lower Jurassic back arc basin; MJU, Middle Jurassic uplift; CCA, Cordilleran continental arc; MSM, Mojave--Sonoran megashear. See text for sources of data.

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their bounding unconformities; (3) Pennsylvanian–Permian, Triassic, and Jurassic sedimentary basins and other areas of accumulation are described and related to tectonic setting; (4) conclusions are presented in the context of tectonic evolution and controls of the region throughout the stratigraphic interval of the chapter. In order to provide a coherent discussion of basin development and geologic history of the region, analysis of each of the above sections is chronological. Emphasis is on a comprehensive treatment of stratigraphy rather than on local or regional details. Although most of the data and many of the conclusions have been presented elsewhere, some of the regional stratigraphy and resulting interpretations are presented here for the first time as a broad regional and stratigraphic synthesis.

2. Precambrian Basement and its Possible Control on Phanerozoic Deposition 2.1. Trends and lineaments The tectonic grain of much of the Colorado Plateau and Southern Rocky Mountains was established during the Proterozoic assembly of southwestern North America. The strong SW-NE grain parallels zones of accretion of the Pinal, Mazatzal, Yavapai, and Mojave provinces (Karlstrom and Humphreys, 1998; Duebendorfer et al., 2006). Long-lasting and prominent lineaments such as the Jemez lineament reflect this grain (Figure 2A). A SENW grain may be related to inherited Precambrian grain as well as wrench tectonics that have affected North America for long periods of time (Stevenson and Baars, 1986). The Colorado Mineral Belt, Uncompahgre, and Zuni lineaments lie on this trend. Both trends and the resulting lineaments have had strong controls on Paleozoic and Mesozoic facies distribution and isopach trends (Blakey, 1988), the distribution of the Ancestral Rockies and associated basins (Stevenson and Baars, 1986; Kluth, 1986), and the Laramide Rocky Mountains (Hoy and Ridgway, 2002).

2.2. Younger Precambrian sedimentary basins During the Middle and Late Proterozoic, a series of sedimentary basins developed across western and southwestern North America. The Apache and Grand Canyon basins developed in the Middle Proterozoic possibly controlled by compressional orogenic events in west Texas and southern New Mexico; the younger Pahrump, Uinta trough, and Belt basins were apparently controlled by the Late Proterozoic rifting of western North America during the breakup of Rodinia (Karlstrom et al., 1999). The effects of these basins on Phanerozoic deposition varied greatly and only a brief review is presented here. The Apache basin was neutral to slightly positive during much of the Paleozoic and Early Mesozoic although the southern margins overlap with the Pennsylvanian–Permian Pedregosa basin. The Grand Canyon basins are mostly on or along the flank of the Kaibab upwarp, a prominent Laramide monoclinal uplift that was generally a positive to neutral area during the Paleozoic. The Uinta trough was an arch during much of the Paleozoic (Ross, 1973) and now forms the core of the Uinta Mountains. The Pahrump and Belt basins evolved into the Cordilleran passive margin during the Paleozoic (Poole et al., 1992).

3. Phanerozoic Tectonics and Depositional History 3.1. Early and Middle Paleozoic Cambrian through Mississippian sedimentation of the southwestern craton was controlled by several major tectonic elements. The Transcontinental arch bounded the southeast margin of the region from the Late Precambrian into the Mississippian (Figure 2A). Early and Middle Paleozoic strata onlap the arch from west to east and Cambrian and Devonian siliciclastic sediments were derived from the positive structure. Cambrian, Devonian, and Mississippian carbonates were deposited in areas not directly affected by clastic sedimentation or during periods of time when clastic input was reduced. The rapidly subsiding Cordilleran miogeocline developed on the passive, western margin of North America during the Late Precambrian and continued throughout much of the Early and Middle Paleozoic (See Miall, Chapter 5, this volume). The eastern margin was the Cordilleran hingeline or Wasatch line (Cordilleran hingeline will be used only during the time when the region was a passive margin; Wasatch Line will be used during subsequent geologic history). Across this curvilinear feature (Figure 2A), nearshore and shallow marine shelf deposits graded westward into thicker shallow marine and

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offshore marine and continental rise-slope deposits (Poole et al., 1992). The resulting sedimentary package, bounded by the above tectonic features, is thin to locally absent east of the hingeline and thickens westward to over 5,000 m in western Utah and eastern Nevada. Cambrian deposits show the most radical thickness changes from east to west. Marine shoreline and local fluvial deposits of the Tapeats Sandstone and related units onlap the Transcontinental arch. Parts of the arch along with other highs in northern and central Arizona were probably never covered by the Tapeats transgressive sequence. The Upper Tapeats grades westward into shallow marine mudstone and sandstone of the Bright Angel Shale and related units; this interval strongly reflects the westward thickening of Cambrian deposits. The Bright Angel grades westward into and is overlain by carbonates of the Muav Limestone. Much of the Muav is shoreline to nearshore marine and probably reflects prolific carbonate deposition that was able to keep up with rising Cambrian sea level. Ordovician and Silurian deposition continued to reflect Late Cambrian carbonate patterns although neither period is present in rocks on the Colorado Plateau and are rare across the Southern Rocky Mountain region. Likely thin deposits blanketed parts of the region but were removed by pre-Devonian erosion. Both systems are represented by extensive carbonate and minor clastic deposition west of the Cordilleran hingeline (Poole et al., 1977). Devonian rocks reflect another transgressive event from west to east onto the Transcontinental arch. Lower Devonian deposits fill sharply defined paleovalleys across the hingeline with mixed carbonate and siliciclastic deposits but are absent eastward. Middle and Upper Devonian carbonates and clastics blanket much of the region but still thin across old paleotopographic highs (Poole et al., 1977). Devonian rocks thicken abruptly west of the Cordilleran hingeline into western Utah and eastern Nevada. During the Late Devonian, a new tectonic setting developed along the western margin of North America; the Antler arc approached the continent and in latest Devonian and Early Mississippian collided with the continent (E. L. Miller et al., 1992; see also Ingersoll, Chapter 11, this volume). The Antler foreland basin formed to the east of the collision zone (Poole et al., 1977). The collision between the Antler arc complex and western North America resulted in thrusting of primarily continental slope-rise deposits over the continental shelf (Burchfiel et al., 1992). The thickened crust caused adjacent foreland basin development. Upper Devonian and Lower Mississippian siliciclastics filled the rapidly subsiding basin; only muds reached the distal portions of the basin along the western margin of the study area. To the east across most of the Western Interior, vast carbonate deposits formed in clear, tropical seas. Carbonates and clastics intertongue along the Wasatch Line region. In response to the Antler foreland basin, a broad forebulge extended along parts of the old Transcontinental arch and Wasatch line as evidenced by onlapping and eastthinning carbonate shoreline deposits (Giles, 1996).

3.2. Pennsylvanian–Permian Pennsylvanian and Permian deposits formed in tectonic settings that contrasted with those of the older Paleozoic. The old Transcontinental arch and passive margin elements were broken by dramatically different tectonic elements (Figure 2B). To the west, the Antler orogenic belt rose along what had been the previous continental margin (E. L. Miller et al., 1992). To the east, the elements of the Ancestral Rockies, sharp uplifts and adjacent basins, extended from eastern Arizona and southern New Mexico northeastward across the cratonic interior to Nebraska (Kluth, 1986). Between these two orogenic zones lies a broad region of cratonic shelf broken by several major basins; these basins were the sites of some of the thickest Phanerozoic sedimentation in the Western Interior. Pennsylvanian and Permian deposits strongly reflect these tectonic elements as well as the cyclic nature of Late Paleozoic global sea level changes (Soreghan, 1994). Siliciclastic sedimentation was strongly controlled by topography. Conglomeratic units flank orogenic highlands and sandstone was spread over adjacent plains by both river systems and eolian dunes. During sea-level lows, much of the region was blanketed by eolian dune deposits (Blakey et. al., 1988). Most continental deposits and some shallow marine and shoreline deposits consist of ubiquitous redbeds. Carbonates were deposited on marine shelves, especially during sea level highs; at times they extended close to orogenic highlands. Basinal deposits include carbonate mudstone, terrigenous mudstone, and locally extensive evaporites. Marine deposits comprise mostly drab tan and gray sedimentary rocks. Shifting shorelines and dynamic tectonic patterns produced complex intercalations of the above deposits, especially evident where red and nonred strata intertongue. During the Late Permian, sea level dropped and deposition ceased across most of the region, even in basinal areas. The resulting Permo–Triassic unconformity extends across the region and although obvious in most places, it can be difficult to locate where Triassic redbeds overlie Permian redbeds.

3.3. Triassic Triassic rocks comprise mostly red continental deposits east of the Wasatch line and gray and tan marine carbonate rocks to the west; strata thicken abruptly west of it (Figure 2C). The thickening was related to a developing

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back-arc basin following the Sonoman orogeny (Lawton, 1994; Ingersoll, Chapter 11, this volume). The Sonoman orogeny involved arc collapse and collision with North America and marked a shift in tectonic patterns (Saleeby and Busby-Spera, 1992; Ingersoll, Chapter 11, this volume). Following the orogeny, the Cordilleran arc developed along much of western North America and a series of back-arc basins formed across the western margin of the craton. A broad platform, mostly mantled with Triassic fluvial deposits, extended eastward onto the craton. Triassic rivers flowed dominantly towards the northwest (Blakey, 1989; Blakey et al., 1993). Elements of the Ancestral Rockies uplands persisted in Colorado and shed some detritus into Triassic rivers, but much of the clastic deposits came from the Appalachian–Ouachita Mountains that marked the Pangaean suture (Blakey, 1994; Riggs et al., 1996; Dickinson and Gehrels, 2003). Lower Triassic rocks thicken westward across the platform and then dramatically thicken across the Wasatch line. Upper Triassic rocks were eroded to the west during uplift of the back-arc basin so their original geometry is difficult to determine. Few Upper Triassic continental rocks are preserved west of the hingeline except in marine basins in western Nevada (Silberling and Roberts, 1962; Saleeby and Busby-Spera, 1992).

3.4. Jurassic The dynamics of Jurassic tectonics of western North America are extremely complex and the subject of much controversy, especially in areas west of the Colorado Plateau. Several phases of the Nevadan orogeny resulted from arc-microcontinent collision with western North America and raised uplands to the west of the Colorado Plateau (Lawton, 1994; Ingersoll, Chapter 11, this volume). A back-arc or foreland basin, depending on how one interprets the cause and geometry of the uplands, rapidly developed along the Wasatch line (Bjerrum and Dorsey, 1995). Generally referred to as the Utah–Idaho trough, the basin began subsiding during the Early Jurassic and then rapidly subsided during the Middle Jurassic; it abruptly slowed during the Late Jurassic (Bjerrum and Dorsey, 1995). Marine deposition dominated the basin center. To the east lay a broad bench that was the site of mostly continental Jurassic deposits, in places dominated by eolian deposition (Blakey et al., 1988). On the western margin of the Colorado Plateau, complex intertonguing of marine and continental deposits marks the transition between the two tectonic settings. The greatest change in thickness once again, occurs across the Wasatch line. The Cordilleran arc was well established during the Jurassic (Figure 2D); south of southern Nevada, the arc was built on continental North America (Andean-style arc) while to the northwest, the arc was built on a complex of accreted terrains, most of which were separated from North America by oceanic crust (Saleeby and Busby-Spera, 1992). By the end of the Middle Jurassic, most of the terrains had fused to North America during the Nevadan orogeny and the arc was Andean in style (Ingersoll and Schweickert, 1986). A major drainage change occurred across the Colorado Plateau in the Middle Jurassic. Triassic and Early Jurassic streams flowed northwest from Pangaean topography to the east; Middle Jurassic streams tapped the arc source to the southwest and flowed to the northeast across the Plateau region (Riggs and Blakey, 1993; Blakey, 1994; Blakey and Parnell, 1991). Northerly winds deflated arid fluvial plains and carried eolian sediment into the arc where eolian sandstone is interbedded with arc volcanics (Busby-Spera, 1988). The last phases of the Nevadan orogeny formed a foreland basin and forebulge-backbulge complex along the western margin of the Colorado Plateau (DeCelles and Currie, 1996). Uplands to the west spawned a major fluvial system that flowed eastward across the Colorado Plateau and Southern Rocky Mountains region and deposited the Morrison Formation (Peterson, 1988b). The Morrison and younger rocks are covered in Chapter 9 (Miall et al., this volume).

3.5. Cretaceous Cretaceous deposition across the region occurred in an extensive foreland basin that lay east of the Wasatch line. Pre-Cretaceous sedimentary rocks were buried from 1–5 km beneath the multiple Cretaceous transgressive– regressive sequences. West of the Wasatch line, thrusting of the Sevier orogeny developed extensive highlands that shed detritus eastward into the foreland basin. Cretaceous sedimentary rocks are dominantly siliciclastic and range from conglomerate and conglomeratic sandstone along the Sevier front in western Utah to sandstone and mudstone across the central and eastern parts of the region. Stratigraphic sequences clearly document six to seven major transgressive–regressive cycles and numerous smaller ones; Maastrictian sedimentation records widespread regressive non-marine deposition (Dyman et al., 1994).

3.6. Cenozoic Near the close of Cretaceous deposition, the region was still near sea level; this would change during the Paleogene as Laramide (Rocky Mountain) tectonics resulted in widespread regional uplift. Uplift was punctuated

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by broad to tight folding, especially monoclinal folding, and sharp, fault-bounded uplift (D. M. Miller et al., 1992). Pennsylvanian through Jurassic rocks that are the focus of this study were stripped from the centers of uplifts but remained mostly buried between and on the flanks of uplifts. In basins of Utah and Colorado, locally thick Paleogene deposits further buried older sedimentary rocks. During the Neogene, integration of the Colorado River system resulted in downcutting and exposure of older sedimentary rocks. The intricate, incised drainage system is responsible for the remarkable exposures of Pennsylvanian through Jurassic rocks for which the region is famous.

4. Pennsylvanian–Middle Jurassic Sequence Stratigraphy 4.1. Introduction The following section describes and broadly interprets the strata that comprise Pennsylvanian through Jurassic sedimentary rocks of the Colorado Plateau and Southern Rocky Mountains Region. The section is divided into sequences of accumulation that are bounded by unconformities (Figures 3 and 4) (Table 1). Where possible, existing terminology is used to describe and name these sequences and their bounding unconformities, although no regional scheme currently exists: Pennsylvanian and Permian sequences (Figure 5) are herein given series names or subdivisions of series (i.e. Morrowan; Lower Leonardian). The definition of the sequences and their boundaries follows established patterns of Ross (1973), Blakey and Knepp (1989), Blakey (1996), and Trexler et al. (2004). Lack of sufficient biostratigraphic data in some places makes it difficult to determine whether the series boundary falls exactly within the stratigraphic interval represented by the unconformity. The ages of the sequences and the bounding unconformities must be considered approximate pending further biostratigraphic studies. Triassic sequences are named after subdivisions of regional formations (i.e. Lower Moenkopi Formation) as established by Blakey et al. (1993) and Blakey and Gubitosa (1983). Jurassic sequences are modified from Pipiringos and O’Sullivan (1978) and Blakey (1994) and are also named after lithostratigraphic units. Note that the method of naming sequences and unconformities varies between these various studies. I favor de-emphasizing

Figure 3 Pennsylvanian--Permian time-stratigraphic diagram. Gray, unconformities; solid horizontal lines, series boundary, presence of unconformity uncertain; ages: left-hand column, compiled byTrexler et al. (2004), right-hand column from Gradstein et al. (2004); stratigraphic data in Cordilleran miogeocline from Snyder (2000, personal communication), other columns compiled from many sources. No vertical or horizontal scale.

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Figure 4 Restored panel cross-sections of Pennsylvanian and Permian rocks, southernWestern Interior. Sections are designed to show stratigraphic relations and general geometry of stratigraphic units; thrust faults along Uncompahgre uplift are diagrammatic. Geometry west of Wasatch Mountains are very diagrammatic. Unconformities shown as heavy lines but not labeled. See Table 1 for formation abbreviations. Lines of section shown on Figure 5.

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List of lithostratigraphic abbreviations used on cross sections.

Jurassic Ja, Aztec Ss; Jcl, lower Carmel Fm; Jcu, upper Carmel Formation; Je, Entrada Ss; c, Cow Springs Mbr; Jk, Kayenta Fm; Jm, Morrison Formation: b, Brushy Basin Mbr; w, Westwater Mbr; r, Recapture Mbr; s, Salt Wash Mbr; Jmo: Moenave Fm; s, Springdale Ss Mbr; d, Dinosaur Canyon Mbr; Jn, Navajo Ss; Jnu, Nugget Ss; Jp, Page Ss; Jpr, Preuss Fm; Jr, Romana Ss; Js, Stump Fm; Jsc, Curtis and Summerville Fms; Jt, Temple Cap Fm; Jtc, Twin Creek Fm; u, upper part; l, lower part; g, Gypsum Springs Mbr; Jw, Wingate Ss; Jwa, Wanakah Fm Triassic Trc–Chinle Formation: c, Church Rock Mbr; o, Owl Rock Mbr; p, Petrified Forest Mbr; mb, Moss Back Mbr; b, Monitor Butte Mbr; s, Shinarump Mbr; a, Ankareh Mbr (or Fm); g, Gartra Mbr Trm–Moenkopi Formation: w, Wupatki Mbr; h, Holbrook and Moqui Mbrs; tp, Timpoweap Mbr; l, Lower Red Mbr; v, Virgin (Ls) Mbr; m, Middle Red Mbr; s, Shnabkaib Mbr; u, Upper Red Mbr; b, Black Dragon Mbr; s, Sinbad (Ls) Mbr; t, Torrey Mbr; m, Moody Canyon Mbr; am, Mahogany Mbr (or mbr Ankareh Fm); Trt, Thaynes Fm; Trdw, Dinwoody Fm; Trw, Woodside Fm Permian Pc, Coconino Ss; Pcl, Colina Ls; Pcm, Cedar Mesa Ss; Pco, Concha Ls; Pct, Cutler Arkose; Pdc, DeChelly Ss; Pdk, Diamond Creek Fm; Pe, Esplanade Ss; Pep, Epitaph Dol; Per, Earp Fm; Pgf, Gerster, Franson fms; Pgr, Grandeur Fm; Ph, Hermit Fm; Pha, Halgaito Fm; Pk, Kaibab Fm; Pkr, Kirkman Fm; Plo, Loray Fm; Ply, Lyons Ss; Pma, Maroon Fm; Poq, Oquirrh Gp; Por, Organ Rock Fm; Pp, Pakoon Ls; Ppl, Plympton Fm; Ppq, Pequop Fm; Pq, Queantoweap Fm; Pr, Rain Valley Fm; Prx, Rex Chert; Ps, Scherrer Fm; Psh, Schnebly Hill Fm; Pt, Toroweap Fm; Pw, White Rim Ss; Pwb, Weber Ss Pennsylvanian IPb, Bird Spring Gp; IPb, Black Prince Ls; IPbm, Beldon and Minturn fms; IPc, Callville Ls; IPer, Earp Fm; IPfo, Fountain Arkose; IPh, Horquilla Ls; IPh, Honaker Trail Fm; IPm, Manakacha Fm; IPma, Maroon Fm; IPp, Paradox Fm; IPpt, Pinkerton Trail Fm; IPw, Wescogame Fm; IPwa, Watahomigi Fm

numerals in sequence boundaries as new work that refines the sequences cannot maintain a parallel system of nomenclature (what can a newly defined unconformity that fits between the J-2 and J-3 be called, for example?). Lithostratigraphy follows established nomenclature across the region although I make no attempt to list every local formational name. This is particularly true for Pennsylvanian and Permian rocks where local partitioning by basins and arches has resulted in a complex array of areally restricted stratigraphic names (Figure 3).

4.2. Pennsylvanian 4.2.1. Mississippian–Pennsylvanian boundary The Mississippian–Pennsylvanian unconformity is everywhere well developed across the region and is generally easy to pick on moderate to poor rock exposures. In many places, Lower Mississippian gray limestone is overlain by Lower to Middle Pennsylvanian red sandstone and mudstone. Rarely can any angularity be detected at outcrop scale. The contact is locally marked by an irregular erosion surface that may be mantled by weathered cherty limestone debris, bedded conglomerate, and red ‘‘tera rosa’’ pedogenic deposits (Blakey and Knepp, 1989). Along many of the uplifts of the Ancestral Rockies, the sub-Pennsylvanian unconformity is cut down through Devonian and Cambrian rocks and rests on Precambrian basement. In some sections, especially in Pennsylvanian basins, Pennsylvanian limestone rests on Mississippian limestone. An erosional surface may or may not be evident. Basal coarse deposits, usually limestone- or chert–pebble conglomerate, locally mark the unconformity. 4.2.2. Morrowan sequence Morrowan strata (Format A of Ross, 1973) are irregularly distributed mostly along the margins of the region (Figure 6). Ross’ use of the term format is synonymous with sequence as used today. East of the Wasatch line, deposits are mostly less than 100 m thick but in western Utah and eastern Nevada, Bissell (1974) reported up to 500 m of Morrowan strata. Lithology ranges from dominantly carbonate rocks with intercalated mudstone in SE Arizona (Ross, 1973; Blakey and Knepp, 1989) and western Utah (Bissell, 1974) to red sandstone and mudstone and thin carbonate in northwestern Arizona (Blakey, 1990) and northern Colorado and Utah (Johnson et al., 1992). Most deposits are believed to be marine, but most lack detailed modern facies analysis. Morrowan strata represent the first incursion of Pennsylvanian marine deposits into the Western Interior following significant hiatus and erosion during the Late Mississippian. The unconformity was of short duration in

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Figure 5 Summary map of Pennsylvanian and Permian basins, uplifts, and other features of the study area. Thickness compiled and averaged from many sources. Blue, basins; light tan, shelves and arches; purple-gray, Pennsylvanian absent, Permian rocks overlie Mississippian rocks; brown, Precambrian basement exposed during Pennsylvanian, part of Permian, and locally into Mesozoic. Thicknesses are given in kilometers.

the negative Cordilleran miogeocline. The geometry of sediment distribution in Arizona, Utah, and Colorado may foreshadow negative areas of Ancestral Rockies tectonics but there is no evidence of uplift this early in the Pennsylvanian (Kluth, 1986).

4.2.3. Unconformity 2 (C-4) Wherever Morrowan strata are present, they are apparently separated from overlying Atokan strata by Unconformity 2. Ross (1973) defined the unconformity in the Pedregosa Basin between the Black Prince

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Figure 6 Morrowan paleotectonics, facies, and isopach map. Dark blue, active basins; brown, incipient Ancestral Rocky Mountains.

Limestone and overlying Horquilla Limestone. In the Grand Canyon, it probably correlates with the unconformity between the Watahomigi and Manakacha formations (McKee, 1982) and Trexler et al. (2004) reported an unconformity at a similar stratigraphic position across north-central Nevada (their C-4). No angularity has been reported and only McKee (1982) has described the surface with any detail. Unconformity 2 (C-4) is probably a sea level low, but detailed regional work would be necessary to confirm such an interpretation.

4.2.4. Atokan sequence Atokan strata (Format B of Ross, 1973) are widely distributed across the region, and are well represented in all Pennsylvanian basins (Figure 7); they are absent from most intervening more positive areas. Strata of this interval contain the oldest significant eolian deposits of the Western Interior (Blakey et al., 1988). Although facies type and distribution are variable and locally complex, a common pattern in several basins is marginal quartz sandstone and carbonate grainstone surrounding micritic carbonate and terrigenous mudstone in basin centers. Patterns of marine and non-marine facies also show a circular distribution with marine deposits dominating basin centers. The above patterns clearly show that many Pennsylvanian basins were well developed during the Atokan (Kluth, 1986). Arkosic conglomerate along the Central Colorado Trough and in NC Colorado documents early phases of Ancestral Rockies uplift (Casey, 1980; Kluth, 1986).

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Figure 7 Atokan paleotectonics, facies, and isopach map. Dark blue, active basins; brown, Ancestral Rocky Mountains (Precambrian-cored uplifts). Queried area is the region where activity on Uncompahgre uplift is uncertain (see text for discussion).

4.2.5. Unconformity 3 (C-5) Unconformity 3 separates Atokan from DesMoinesian rocks across the study region (Ross, 1973). The unconformity is mainly determined by paleontological studies and can be difficult to pick on outcrop. Defined in the Pedregosa Basin, it probably correlates with unconformity (C-5) at a similar stratigraphic positions in Nevada (Trexler et al., 2004). 4.2.6. Desmoinesian sequence Desmoinesian strata (Format C-G of Ross, 1973) are widely distributed across the region. In the Pedregosa Basin, Ross actually recognized five formats and intervening unconformities but they are difficult to separate without good fossil control and their regional correlation as individual formats has never been accomplished; Blakey and Knepp (1989) lumped the five formats and that practice is followed here. Present distribution and facies patterns are strongly controlled by the geometry of Pennsylvanian basins (Figure 8). Each basin has at least 300 m of Desmoinesian strata and the Central Colorado Trough, Paradox Basin, and Oquirrh Basin each contain in excess of 1,500 m of strata. All basins including the Cordilleran basins display well-developed marginal platform and shelf deposits that wrap around basinal deposits. Marine shelf deposits are dominated by fossiliferous carbonate grainstone and packstone and basin centers contain dark carbonate mudstone and terrigenous mudstone (Johnson

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Figure 8 Desmoinesian paleotectonics, facies, and isopach map. Dark blue, active basins; brown, Ancestral Rocky Mountains; purple, evaporite deposition. Queried area is the region where activity on Uncompahgre uplift is uncertain; connection between Paradox basin and Eagle basin uncertain (see text for discussion).

et al., 1992). The Central Colorado Trough and eastern Paradox Basin are rimmed by thick, coarse, arkosic conglomerate that fines abruptly into basin platforms and centers (Casey, 1980; Baars and Stevenson, 1981; Johnson et al., 1992; Karachewski, 1992). Both basins also contain extensive marine evaporite deposits (Wengard and Matheny, 1958; Hite, 1970; Tweto, 1977). Figure 8 queries the existence of the Uncompahgre uplift at this time in WC Colorado. Chuck Kluth (personal communication, 2007) has unpublished seismic data coupled with well data that suggest that the northern Central Colorado trough (Eagle basin) and Paradox basin were one entity during the Desmoinesian and Missourian. Desmoinesian strata were deposited in basins across the southern Western Interior during the first (main) pulse of the Ancestral Rockies (Kluth, 1986). The flanks of basins adjacent to uplifts reflect strong influx of immature, coarse, arkosic sediments. Sharp uplift flanked by equally strong subsidence and sediment thickness, locally in excess of 1,500 m in some basins, attests to this.

4.2.7. Unconformity 8 (C-6) Unconformity 8 was defined by Ross (1973) in eastern Arizona where the surface separates Desmoinesian from Missourian strata. The unconformity may correlate with a pre-Missourian unconformity in the Paradox Basin

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recognized by Welch (1958) and unconformity C-6 in NE Nevada and adjacent Utah (Trexler et al., 2004). The unconformity may be coincident with reduced rates of uplift in the Ancestral Rockies and corresponding lowered rates of subsidence in basins as reported by Kluth (1986).

4.2.8. Missourian sequence Missourian rocks (Format H-I of Ross, 1973) are similar in distribution and lithology to those of underlying Desmoinesian rocks. In many basins Missourian rocks show an increase in siliciclastic sedimentation at the expense of carbonate deposition and are commonly banded, cyclic deposits of red, tan, and gray mudstone, sandstone, and carbonate, respectively (Ross, 1972; Blakey and Knepp, 1989). Kluth (1986) recognized that vigorous uplift of the Ancestral Rockies continued into Missourian time but at a reduced rate compared to that of Desmoinesian time. Basin subsidence also was reduced (Figure 9). It must be cautioned that comparing rates of sedimentation and subsidence between epochs and sequences should be done with care as duration of different sequences is rarely equal; note also that the entire Middle Pennsylvanian (Desmoinesian–Missourian) is only 5 Myr duration on the 2004 timescale (Gradstein et al., 2004).

Figure 9 Missourian paleotectonics, facies, and isopach map. Dark blue, active basins; brown, Ancestral Rocky Mountains; queried area is the region where activity on Uncompahgre uplift is uncertain.

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4.2.9. Unconformity 10 Unconformity 10 separates Missourian from Virgilian rocks in SE Arizona (Ross, 1973). The unconformity may represent a short period of time, although rocks above it commonly show a sharp increase in sandstone and mudstone (Ross, 1973; Blakey and Knepp, 1989). 4.2.10. Virgilian sequence Virgilian rocks (Format J-L of Ross, 1973) comprise widespread siliciclastic deposits across much of the region (Figure 10). Although carbonate deposition continued in some basin centers and on shelves removed from uplifts, marine, eolian, and fluvial sandstone and associated red mudstone dominate most areas (Loope, 1984; Blakey and Knepp, 1989). Kluth (1986) reported a general continued slowdown of uplift and subsidence during Virgilian time although there remained local areas of vigorous tectonic activity.

4.3. Permian 4.3.1. Pennsylvanian–Permian boundary The boundary between Pennsylvanian and Permian rocks across the region has generally been recognized as an unconformity, although in many places, the contact is difficult to pick on outcrop (Peirce, 1989). This difficulty

Figure 10 Virgilian paleotectonics, facies, and isopach map. Dark blue, active basins; brown, Ancestral Rocky Mountains; queried area is region where separation may have existed between Uncompahgre and San Luis uplifts (see text for discussion).

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exists because of similarity of youngest Pennsylvanian strata with those of the overlying Permian; in many places redbeds rest on redbeds and in other places local conglomerate marks the disconformity (Blakey and Knepp, 1989).

4.3.2. Wolfcampian sequence Blakey (1996) recognized four Permian sequences and their intervening unconformities across the Colorado Plateau and these Permian elements are used in this chapter (Figure 3). Wolfcampian strata (Sequence P-1 of Blakey, 1996) are widespread across the Colorado Plateau and Southern Rocky Mountain region and are present most places except across the highest Ancestral Rockies uplifts (Figure 11). Wolfcampian strata comprise mixed lithologies that reflect local tectonic and topographic conditions. Thick arkosic deposits are present in the eastern Paradox Basin (Campbell, 1980; Mack and Rasmussen, 1984) and thinner arkoses flank other basins. Eolian sandstone is widespread, especially across the Colorado Plateau (Baars, 1962; Blakey, 1996). Carbonates interfinger with both lithologies and dominate deposition in western Utah and adjacent Nevada (Bissell, 1970). The various sedimentary facies accumulated in mixed marine, fluvial, and eolian environments (Blakey, 1996; Blakey, 2002; Langford and Chan, 1989).

Figure 11 Wolfcampian paleotectonics, facies, and isopach map. Dark blue, active basins; brown, Ancestral Rocky Mountains; queried area is the region where separation may have existed between active Uncompahgre and waning San Luis uplifts (see text for discussion). Gray areas where Lower Permian is absent were likely waning uplifts.

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In many basins of the Ancestral Rockies, especially the Central Colorado trough, Permian deposits are markedly thinner than underlying Pennsylvanian deposits. This suggests a general slowdown in the Late Paleozoic orogeny (Kluth, 1986; Miall, Chapter 8, this volume); however, rates of sedimentation in parts of the Paradox Basin and in areas of western Utah and eastern Nevada show a continued rapid rate of subsidence and sedimentation (Bissell, 1970; Blakey, 1996). Clearly the Ancestral Rockies orogeny was not finished by Early Permian time.

4.3.3. Unconformity P-sc Unconformity P-sc (Blakey, 1996) marks a break to rapid Wolfcampian sedimentation. In parts of western and northern Colorado, Wolfcampian rocks are the youngest Paleozoic deposits. Elsewhere, Wolfcampian rocks are overlain by unconformity P-sc. The unconformity is planar in many areas and is only confirmed by regional stratigraphic relations (Blakey, 1996). The unconformity may correlate with unconformity P-2 in the Cordilleran miogeocline (Trexler et al., 2004). The hiatus may lie within the Wolfcampian or Leonardian or mark the boundary between the two.

4.3.4. Lower Leonardian sequence Lower Leonardian rocks (Sequence P-2 of Blakey, 1996) are unevenly distributed across the study area (Figure 12). In the Denver, Orogrande, and Holbrook basins, sandstone and red mudstone of eolian, restricted marine, and sabkha origin dominate (Blakey et al., 1988). Thicknesses in the Holbrook and Orogrande basins exceed 600 m (Blakey, 1990) In western Utah and Nevada, carbonates are intercalated with quartz sandstone and thickness exceeds 1,000 m (Bissell, 1970). Lower Leonardian rocks are apparently absent across much of the central portion of the region. In the southern Colorado Plateau region, the top of the sequence comprises the widespread Coconino–Glorieta eolian complex, probably the largest Late Paleozoic eolian deposit of the Western Interior (Blakey et al., 1988). The Lower Leonardian Sequence may reflect changing tectonic conditions across the region. Older Pennsylvanian and Permian sequences clearly reflect Ancestral Rockies tectonics (Kluth, 1986). Armin (1987) suggested that the Holbrook and Orogrande basins might be backbulge deposits related to thrust loading in the Marathon thrust belt of west Texas. The deep subsidence in the Cordilleran miogeocline could be related to tectonic events in the Cordilleran region (Trexler et al., 2004).

4.3.5. Unconformity P-tw An unconformity within the Leonardian extends across the region (sub-Toroweap of McKee, 1938; P-tw of Blakey, 1996; P-3 of Trexler et al., 2004). The unconformity is subtle on many outcrops, especially east of the margin of the Toroweap Formation, and is most easily seen by regional stratigraphic patterns (Blakey, 1996). Across the southern and western Colorado Plateau, the unconformity is marked by the marine-flooding surface of the Toroweap marine transgression.

4.3.6. Upper Leonardian sequence The Upper Leonardian Sequence (Sequence P-3 of Blakey, 1996) comprises marine deposits of the Toroweap Formation and coeval eolian facies (Chan, 1989). Thin evaporites are present in the Grand Canyon region (Rawson and Turner-Peterson, 1980). The sequence is absent in the Four Corners region and across most of western Colorado (Figures 13 and 14). It is generally less than 100 m thick across the study area; however, it thickens to over 1,500 m in parts of the Cordilleran miogeocline (Bissell, 1970). On the Colorado Plateau, the sequence was deposited during a general marine transgression–regression cycle (Rawson and Turner-Peterson, 1980). The sequence apparently lacks coarse material derived from Ancestral Rockies uplifts suggesting that the mountains were worn down by this time.

4.3.7. Unconformity P-k Unconformity P-k separates the Kaibab Formation from the underlying Toroweap Formation across the western and southern Colorado Plateau (McKee, 1938). In the Cordilleran basin, Trexler et al. (2004) recognized the P-4 unconformity at a similar horizon.

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Figure 12 Lower Leonardian paleotectonics, facies, and isopach map. Gray areas where rocks are absent were likely waning uplifts or remnant uplands; queried areas are where Lower Leonardian rocks are probably absent, ruled brown area is where Lower Leonardian rocks cannot be separated fromWolfcampian rocks.

4.3.8. Guadalupian sequence and younger Permian rocks The Guadalupian Sequence (Sequence P-4 of Blakey, 1996) is present mainly across the western and southeastern portions of the study area where the Kaibab and San Andres formations were deposited (Figure 13). The sequence likely contains youngest Leonardian as well as Guadalupian rocks (McKee, 1938). Limestone and sandy dolomite of shallow marine and shoreline origin dominate the sequence (Irwin, 1971). The sequence is generally less than 200 m thick except in the Cordilleran miogeocline where it thickens to over 1,000 m (Bissell, 1970). The sequence was deposited during the last major Paleozoic marine transgression–regression cycle across the southern Western Interior (Blakey, 1996). The Kaibab Formation (and related rocks) is the youngest Permian unit across most of the study area. West of the Wasatch Line, and in the northern extremities of the study area, younger Permian rocks were deposited in several marine transgressive–regressive cycles (Whalen, 1996).

4.4. Triassic 4.4.1. Introduction Triassic rocks are generally well exposed across the Colorado Plateau, but less well exposed in the Southern Rocky Mountains. Most Lower and Middle Triassic rocks are assigned to the Moenkopi Formation and most

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Figure 13 Latest Leonardian and younger (mostly Guadalupian--Kaibab Fm and post-Kaibab) paleotectonics, facies, and isopach map. Dark blue, active basins.

Upper Triassic rocks are assigned to the Chinle Formation (Figures 15 and 16). Lower Triassic rocks have similar geometry, dramatic thickening west of the Wasatch line (Figure 17), to that of many Paleozoic systems. Upper Triassic rocks display contrasting geometry and thicken to the SE and NE (Figure 18).

4.4.2. Unconformity Tr-1 In most places of the southern Western Interior where Triassic rocks rest on Permian rocks, they overlie an obvious unconformity based on both stratigraphic and paleontologic evidence (Stewart et al., 1972b). At the north edge of the study area, the boundary is less certain and locally suggests conformable relations between Permian and Triassic rocks (Stewart et al., 1972b). In the San Rafael Swell, thin limy sandstone deposits lie between undoubted Permian rocks below and Triassic rocks above and contain a molluscan fauna that may be either Permian or Triassic in age (Blakey, 1974). South of these two areas, Unconformity Tr-1 generally separates rocks of Leonardian or Guadalupian age from Lower or Middle Triassic rocks above (Pipiringos and O’Sullivan, 1978; Blakey et al., 1993). The unconformity developed during the Permo–Triassic sea level low stand, one of the lowest in the stratigraphic record (Haq et al., 1988; Haq, 1991). Local relief on the surface approaches as much as 100 m in NW Arizona and SW Utah (Nielson and Johnson, 1979). Except for local angular unconformities in the Salt Anticline region, no angular discordance has been reported (Blakey, 1974).

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Figure 14 Columnar columns and correlation of Permian rocks on the central and western Colorado Plateau. Sequences and sequence boundaries used in this report are shown. These rocks are responsible for much of the rugged scenery of numerous national parks, monuments, and wilderness areas of the region. See Figure 4 for key to lithology.

4.4.3. Lower Moenkopi sequence The oldest Mesozoic sequence in the study area comprises the lower part of the Moenkopi Formation (Blakey et al., 1993); it was deposited across parts of northern Arizona and much of Utah (Figure 19). The sequence consists of deposits less than 100 m thick along the eastern margin but thickens westward to over 2,000 m in NW Utah. As defined herein, the Lower Moenkopi sequence includes strata that lie below the lower massive sandstone of McKee (1954) and coeval Shnabkaib Member; this horizon is coincident or very close to a regional magnetostratigraphic boundary in the Lower Triassic; fossils date the sequence as Lower Triassic (Morales, 1987). The Lower Moenkopi sequence comprises distinct eastern and western lithofacies with an intermediate zone where the two are intercalated. The eastern facies consists of pale to dark reddish-brown mudstone and very finegrained sandstone. Bedding is generally thin and rhythmic and forms ledgy slopes. Ripple marks, ripple cross lamination and mudcracks are dominant but a wide range of sedimentary structures indicative of shallow water deposition are present (Blakey, 1974; Dubiel, 1994). The western facies is dominantly gray to tan limestone. Bedding is generally thin and rhythmic, although more massive-weathering units are present. A wide range of carbonate fabric and textures and locally abundant fossils document shallow marine shoreline to offshore marine deposition (Paull and Paull, 1993). The eastern and westen facies are complexly intercalated across a zone that trends NNE-SSW across central and western Utah (Blakey et al., 1993). Carbonate percentage increases westward across the zone and thin carbonate marker beds penetrate eastward into the red sandstone and mudstone. The Lower Moenkopi sequence was deposited on a broad, flat, relatively featureless coastal plain (Blakey, 1974; Ochs and Chan, 1990). The coastal plain was so flat that slight changes in relative sea level caused widespread marine transgression and regression. Fine siliciclastic material was trapped and reworked by low

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Figure 15 Triassic--Jurassic time-stratigraphic diagram. Gray, unconformities. Ages: left-hand column, Harland et al. (1989); right-hand column, Gradstein et al. (2004). Numbered unconformities from Pipiringos and O’Sullivan (1978); J-sk and J-sup from Blakey (1994). No vertical or horizontal scale.

energy coastal environments, especially broad intertidal flats. Carbonate environments persisted in clearer shallow marine environments, especially during weak transgressive episodes while sand and mud were trapped on the coastal plain (Blakey et al., 1993). The Lower Moenkopi sequence represents rather mild tectonic events across the Western Interior. Shoreline position, depositional environment, and carbonate versus siliciclastic deposition were controlled by sediment influx from the east and relative sea level during a time of steady subsidence with subsidence rates dramatically increased to the west (Blakey et al., 1993). 4.4.4. Unconformity Tr-lm Unconformity Tr-lm (under the lower massive sandstone) is a widespread but very subtle planar surface that separates similar lithology above and below. The unconformity is difficult to locate on local outcrops but is apparent based on regional trends. The surface overlies strata that dramatically thicken to the west and is overlain by strata that maintain a rather constant thickness across much of the region (Blakey et al., 1993). 4.4.5. Upper Moenkopi sequence The Upper Moenkopi sequence consists of red sandstone and mudstone that forms ledges and weak cliffs where present. The sequence is presently recognized across much of northern New Mexico and Arizona and southwestern Utah and parts of adjacent Nevada (Figure 20). It may be present elsewhere farther north, although it is clearly removed by pre-Chinle erosion in many areas (Blakey, 1974). Unlike the Lower Moenkopi sequence, the Upper Moenkopi sequence does not display a marked westward thicking. Fluvial channel deposits and adjacent flood plain deposits dominate the sequence; the southern and eastern areas of deposition are characterized by perennial meandering stream deposits that grade northward (distally) into ephemeral stream deposits. In NW Arizona and SW Utah, at least part of the sequence includes mudstone and gypsum coastal plain deposits of the Shnabkaib Member (Blakey et al., 1993). The sequence is Middle Triassic based on a widespread vertebrate fauna (Morales, 1987).

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Figure 16 Triassic restored panel cross-sections, southernWestern Interior. Sequences and sequence boundaries are shown. See Table 1 for formation abbreviations. Purpose of these sections is to show regional relations rather than local details.

The Upper Moenkopi sequence marks a change in subsidence history of the region. Although streams clearly flowed northwesterly across the region, similar to stream flow below, the lack of northwest thickening suggests that rapid subsidence along and west of the Wasatch Line had ended. The sequence may mark a transition between the contrasting isopach patterns of the Lower Moenkopi Formation and overlying Chinle Formation (compare Figures 17, 18, and 20).

4.4.6. Unconformity Tr-3 Unconformity Tr-3 is one of the most dramatic Triassic–Jurassic unconforities across the Western Interior (Pipiringos and O’Sullivan, 1978). It separates with obvious disconformity the Moenkopi and Chinle formations, except where the former is absent; where the Moenkopi is absent the Chinle overlies rocks as old as Precambrian (Pipiringos and O’Sullivan, 1978). On a regional scale, the unconformity is locally angular, especially near elements of the Ancestral Rockies, but the angularity is so slight as to not be observable on outcrop. Relief on the unconformity is obvious, both at local and regional scales. Documented relief approaches 100 m and 10–20 m is apparent on many outcrops (Blakey, 1974; Blakey et al., 1993). The unconformity represents an obvious change in base level and a change in fluvial regime and style. Both Moenkopi and Chinle streams had NW to N paleoflows but here the similarities end; Moenkopi streams flowed in arid climates and were at least partly ephemeral; except very locally, they lacked any siliceous extrabasinal conglomerate and flowed near base level on coastal plains. Chinle streams flowed in humid to semiarid climates and were mostly perennial; they contained abundant siliceous extrabasinal conglomerate and show evidence of complex and changing base level (Kraus and Middleton, 1987). How these contrasts between fluvial systems were orchestrated by tectonic, eustatic (or other base level changes), or climatic controls has yet to be determined in detail.

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Figure 17 Summary of Lower and Middle Triassic thickness and sub-Triassic paleogeology. East of zero isopach, paleogeology is under Upper Triassic Chinle Formation. Note general paleo-anticlinal SW-NE trend through Ancestral Rockies. Paleogeology modi¢ed from Pipiringos and O’Sullivan (1978).

4.4.7. Lower Chinle formation The Chinle Formation is one of the most thoroughly studied units on the Colorado Plateau. The formation was deposited in complex and varying fluvial systems and related continental environments; most of the rivers flowed towards the north to west (Stewart et al., 1972a; Blakey and Gubitosa, 1983; Dubiel, 1989). The Chinle was deposited under more humid conditions than that of the underlying Moenkopi, possibly under a monsoonal regime (Dubiel et al., 1991). Most recent work on the Chinle Formation suggests that this complex fluvial deposit was the result of two or more sequences of deposition with intervening unconformities. However, there is lack of agreement as to where sequences and sequence boundaries are located within the formation (cf. Lucas and Marzolf, 1993). Based on regional observations, I recognize three sequences within the Chinle Formation and two internal, regional unconformities: one at the base of the Moss Back and Sonsela members and one at the base of the Church Rock and Rock Point Members. Lucas (1993) has proposed a radically different nomenclature system and stratigraphic framework that are not followed here. The Lower Chinle Formation as herein defined includes strata between the basal Chinle unconformity (Tr-3) and a prominent change in sedimentation at a probable unconformity below the Moss Back and Sonsela members. Using this definition of the interval, it includes the Shinarump, Monitor Butte, and lower Petrified Forest members. The interval is generally present across the southern 2/3 of the study area and generally thickens

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Figure 18 Triassic.

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Summary of Upper Triassic thickness. Note how thickness trends contrast markedly with those of Lower and Middle

to the south (Figure 21). It was deposited by NW to N flowing streams and associated flood basins, lakes, swamps, and pedogenic envoronments (Blakey and Gubitosa, 1983; Dubiel, 1994). Sandstone is mostly coarse- to very coarse-grained and commonly is conglomeratic with clasts ranging into the cobble size. The mostly siliceous clasts are variable mixes of quartz, quartzite, and chert, with highly variable (both in composition and percentage) volcanic clasts (Figure 21). Quartz and quartzite are from Precambrian sources, chert is chiefly from Late Paleozoic limestone, and volcanics are from Triassic sources (Stewart et al., 1972a; Stewart et al., 1986). Zircons in some sandstone units document that some streams were sourced by Pangaean highlands (Riggs et al., 1996). Details of the Chinle fluvial systems are well documented (references cited above and references therein).

4.4.8. Unconformity Tr-sm The base of the Sonsela Sandstone Member and likely coeval Moss Back Member is a regional unconformity (Blakey and Gubitosa, 1983). The scoured base of this large fluvial complex displays several meters of local relief and likely tens of meters of regional relief. The unconformity likely represents a change in fluvial style accompanied by regional base level change. Lupe and Silberling (1985) have related this and other Chinle base level changes to Triassic eustatic events based on tentative correlation of the Chinle to marine rocks in western Nevada. The unconformity may mark the Carnian–Norian stage boundary (Lockley and Hunt, 1994).

4.4.9. Middle Chinle formation The middle part of the Chinle Formation (Figure 22) as herein defined includes the Moss Back, Sonsela, Upper Petrified Forest, and Owl Rock members. The general lithology and sedimentary history are similar to those of the Lower Chinle Formation, although important differences exist. The Middle Chinle has a persistent carbonate

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Figure 19 Moenkopi Formation lower marine and £uvial sequence isopach and facies map. Note that facies trends closely parallel westward thickening of interval.

interval, the Owl Rock Member, which was formed in lacustrine and pedogenic settings (Blodgett, 1988; Dubiel, 1994) and has a higher percentage of intrabasinal carbonate clasts that range into the pebble to cobble range (Blakey and Gubitosa, 1983). This limestone–pebble conglomerate is an important host lithology for Chinle uranium ore deposits. Bentonitic mudstone, an important Chinle lithofacies, decreases from south to north across the region. The lithology and depositional environments of the Middle Chinle Formation are well documented in the literature (Blakey and Gubitosa, 1983, 1984; Dubiel, 1989; Dubiel, 1994). 4.4.10. Unconformity Tr-cr Unconformity Tr-cr is marked by a widespread erosional disconformity across the region and a change in fluvial style above. In most aspects, it is similar to Tr-sm. The aspects vary from erosional and irregular to the north to more planar to the south. 4.4.11. Upper Chinle formation The Upper Chinle Formation (Figure 23) as herein defined includes the Church Rock Member to the north and coeval Rock Point Member to the south. This is the reddest interval in the Chinle and likely represents an overall

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Figure 20 Moenkopi Formation upper £uvial sequence isopach and facies map. Interval comprises sheet-like deposits generally less than 100 m thick. Note contrast with patterns of lower marine and £uvial sequence.

increase in aridity (Dubiel, 1994). It is also the least understood portion of the Chinle Formation. In much of Utah and Colorado, the interval is marked by ledge-forming sandstone and conglomerate that fine upwards into red mudstone (O’Sullivan, 1970). Both intrabasinal and extrabasinal conglomerate are present. A complex of braided and meandering streams deposited the northern portion of the interval (Blakey and Gubitosa, 1983; Hazel, 1994). The southern portion of the interval is marked by an increase in tabular eolian and sabkha sandstone and silty sandstone (Dubiel, 1994). Once thought to intertongue with the overlying Wingate Sandstone (Harshbarger et al., 1957), Nation (1990) has documented a low-angle unconformity that separates similar facies of the two formations.

4.5. Jurassic 4.5.1. Introduction Jurassic strata are widely exposed across the study area and form the largest areal outcrop of any geologic system across the Colorado Plateau. Because of the areal continuity of Jurassic rocks on outcrop, nomenclature remains relatively consistent across the region (Figure 15). Jurassic rocks thicken westward into the Utah–Idaho trough (Figure 24). Lower Jurassic rocks (Figure 25) are beveled eastward across the region by the J-2 unconformity

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Figure 21 Chinle Formation lower £uvial system isopach, £uvial trend, and paleovalley map (isopachs modi¢ed from Stewart et al., 1972a). The tan areas de¢ne the broad paleovalleys and were constructed by combining distribution of Shinarump Member, paleocurrent trends, and conglomerate petrology (modi¢ed from Blakey and Gubitosa, 1983). Mudstone and sandstone of the Monitor Butte Member and related units overlie the paleovalleys and covered the region shown in red brown.

(Pipiringos and O’Sullivan, 1978). Middle Jurassic rocks overlap the J-2 surface and extend eastward of the study region (Figure 26).

4.5.2. Unconformity J-0 Unconformity J-0 is a regional unconformity that bevels underlying strata to the southwest (Pipiringos and O’Sullivan, 1978; Blakey, 1994). In SW Arizona and adjacent California, the unconformity truncates Paleozoic strata (Reynolds et al., 1989). Across the study area, the unconformity lacks measurable relief and locally superimposes similar facies making recognition difficult. In spite of obvious regional truncation and tilting associated with the surface (I am unaware of any locally obvious angular unconformity except in SE California and SW Arizona), stream flow was to the NW above and below the unconformity (Blakey, 1994). The unconformity is believed to mark the Triassic–Jurassic boundary, although some paleontological evidence remains highly equivocal (Lockley and Hunt, 1994).

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Figure 22 Chinle Formation middle £uvial system isopach, £uvial trend, and petrographic facies map (isopachs modi¢ed from Stewart et al., 1972a). The correlation of the Sonsela, Moss Back, and Gartra £uvial systems is equivocal. Each is sharply incised into underlying mudstone and ¢nes upward into overlying mudstone and sandstone. The carbonate trend consists of £uvial channel deposits with abundant carbonate clasts that were derived from adjacent caliche overbank deposits and lacustrine systems (Blakey and Gubitosa, 1983).

4.5.3. Lower Glen Canyon Group The Glen Canyon Group represents one of Earth’s greatest fossil desert accumulations. This vast complex of rock was deposited in great interior ergs and by both ephemeral and perennial stream complexes; both erg sequences contain huge eolian complexes to the NE part of the study area bordered by fluvial complexes to the SW. The area of intertonguing between fluvial and eolian systems is well exposed for both sequences and each occupies generally the same area (Blakey et al., 1988). The Lower Glen Canyon Group includes the Wingate Sandstone and coeval Dinosaur Canyon Member of the Moenave Formation. Eolian deposits of the Wingate intertongue across a 150 km wide NW trending band with ephemeral fluvial deposits of the Dinosaur Canyon (Blakey et al., 1988; Clemmensen et al., 1989; Blakey, 1994). Together they form a general sheet-like deposit that gradually thickens and thins without obvious trend across the region (Figure 27). Facies patterns, NW stream flow, SE migrating dunes, and patterns of intertonguing suggest a complex fluvial–eolian recycling between the two units (Blakey, 1994).

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Figure 23 Chinle Formation in upper £uvial system isopach and facies map (isopachs modi¢ed from Stewart et al., 1972a). The Black Ledge £uvial system is sharply incised into underlying mudstone and local sandstone and ¢nes upward into mudstone and sandstone.

4.5.4. Unconformity J-sk Unconformity J-sk is a widespread disconfomity at the base of the fluvial Kayenta Formation (Nation, 1990; Blakey, 1994). The surface marks an obvious break between eolian deposits below and fluvial deposits above. Local relief of up to 10 m is present where fluvial channels incise cliff-forming eolian sandstone. There is no discernable pattern of truncation or tilting below the surface, which suggests climatic and possibly base level change as the cause. The unconformity formed during the Early Jurassic but is difficult to date more precisely.

4.5.5. Upper Glen Canyon Group The Upper Glen Canyon Group comprises the Kayenta Formation and Navajo Sandstone. The Navajo Sandstone may be the largest eolian system in the geologic record (Blakey, 1994). The plan view pattern of intertonguing between parts of the Navajo and Kayenta is similar to that of the underlying Wingate–Dinosaur Canyon (Blakey, 1994); however, the cross sectional geometry of the two is considerably different (Figure 24). Whereas the Lower Glen Canyon Group has an irregular sheet geometry, the Upper Glen Canyon group thickens drastically to the SW (Figure 28). Because the Upper Glen Canyon Group is everywhere overlain by an erosional surface, it is unknown whether preserved isopach patterns of the sequence reflect primarily westerly increase in

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Figure 24 Restored panel cross-sections of Jurassic rocks of southernWestern Interior. Sequences and unconformities are shown. See Table 1 for formation abbreviations. Purpose of these sections is to show regional relations rather than local details.

subsidence rate, uplift and truncation to the east, or a combination of both (Blakey, 1994). The sequence was deposited during a low eustatic interval (Haq, 1991) and it seems likely that moderate subsidence rates across much of the western portion of the study area may have created a topographic basin below sea level, but without direct connection to the sea. Coeval Kayenta fluvial deposits drained into the basin, the beginnings of the Utah– Idaho Trough, and were then deflated by NW winds to partially feed the Navajo ergs (Blakey, 1994; see his Figure 4). The topographic hole coupled with low sea level isolated the region from all other depositional systems for millions of years allowing the Navajo ergs to fill the basin, almost by default. Water from the Kayenta river seeped into the porous sands and maintained a near surface water table during much of Navajo deposition; the high water table trapped eolian sand in the basin preventing excessive loss of sediment. The high water table is well documented by widespread fresh water carbonates and mass flow sand deposits (Eisenberg, 2003, and references cited therein). Dickinson and Gehrels (2003) demonstrated a dominantly cratonic North American and Pangaean source for the Navajo and other Colorado Plateau erg systems based on zircon ages within eolian sandstone. 4.5.6. Unconformity J-1 The J-1 unconformity is restricted by subsequent erosional patterns to the western margin of the study area in SW Utah (Figure 29). The surface truncates the eolian Navajo Sandstone and is succeeded by sabkha deposits of the Temple Cap Sandstone (Peterson and Pipiringos, 1979). On most outcrops, the surface is planar, although rarely is it cleanly exposed. Kowallis et al. (2001) established a minimum age of 170 Ma for the unconformity based on 39Ar/40Ar ages from intercalated bentonites in the overlying Temple Cap Sandstone. The restricted exposures in the southern Western Interior offer few clues as to the origin of the J-1 surface; however, north of the study area, the J-1 is attributed to a Middle Jurassic lowstand (Brenner and Peterson, 1994). 4.5.7. Temple Cap Sandstone The Temple Cap Sandstone is restricted to SW Utah and consists of two facies, red sandy mudstone of sabkha origin and tan cross-bedded sandstone of eolian origin (Peterson, 1994). Based on stratigraphic position, the unit

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Figure 25 Glen Canyon Group summary isopach map. Modi¢ed from Blakey et al. (1988).

is early Middle Jurassic (Aalenian) and is related to the first Middle Jurassic transgression into the Western Interior (Peterson and Pipiringos, 1979; Peterson, 1994). The remnant outcrops in SW Utah probably bordered a restricted Jurassic seaway that lay to the NW (Figure 29). 4.5.8. Unconformity J-2 The J-2 unconformity is one of the most profound Mesozoic unconformities (Figure 24) and marks the boundary between the Absaroka and Zuni Sequences. Sloss (1988) speculated that the formation of the J-2 surface was the major controlling event of the current geologic outcrop patterns of central North America. The J-2 can be correlated across the entire Western Interior and into adjacent regions (Pipiringos and O’Sullivan, 1978). Across the study area, the surface cuts down through older strata, resting on Precambrian rocks across elements of the Ancestral Rockies. Much of the eastward thinning of the Navajo Sandstone may be due to J-2 erosion, although this cannot be determined with certainty until internal correlation of the Navajo is established (Blakey, 1994). Although in negative areas of the northern Western Interior the unconformity only represents 2–3 Myr, in most regions more time is represented; the complex paleogeology below the surface must clearly be the result of tectonic warping of much of the craton (Sloss, 1988). Across the study area, the J-2 surface where cleanly exposed always documents at least several meters of local relief to local extremes of nearly 50 m in south-central Utah (Blakey et al., 1996).

Pennsylvanian–Jurassic Sedimentary Basins of the Colorado Plateau

Figure 26

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San Rafael Group summary isopach map. Compiled from many sources.

4.5.9. Page-Lower Carmel The Page Sandstone and coeval parts of the Lower Carmel Formation document complex interactions between coastal eolian systems and interior restricted marine systems (Blakey et al., 1996; Havholm et al., 1993). The sequence was deposited in and along the eastern margin of the rapidly subsiding Utah–Idaho trough (Blakey et al., 1996; Peterson, 1994) which may reflect the initial stage of Mesozoic foreland basin development across the Western Interior (Bjerrum and Dorsey, 1995). Details of the sedimentary history, relations between marine and eolian events, and eustatic and tectonic controls on deposition were documented by Blakey et al. (1996). A normal marine fauna in the Lower Carmel Formation is Bajocian age (Imlay, 1967, 1980). The interval thickens from a depositional margin across SE Utah to over 200 m in SW Utah; eolian strata to the east are intercalated with and are replaced westward by sabkha and restricted shallow marine redbeds and marine limestone (Figure 30). The interval contains abundant volcanic material, ash beds and volcanic grains, which reflect activity in the Cordilleran Arc to the SW (Blakey and Parnell, 1991).

4.5.10. Unconformity J-sup A prominent regional unconformity, J-sup (upper surface on the Page Sandstone), separates the lower and upper portions of the Carmel Formation and related rocks (Blakey et al., 1996). The resulting surface is planar to slightly

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Figure 27 Lower Glen Canyon Group isopach and £uvial--eolian paleo£ow map (modi¢ed from Blakey et al., 1988). Note strongly opposed £uvial and eolian paleocurrents.

undulatory and in many places is overlain by a prominent bentonite. In some areas, the surface separates eolian sandstone of the Page Sandstone below from red sandstone and mudstone of the Upper Carmel Formation above. Detailed stratigraphic studies demonstrate that the surface cuts down section to the NE with removal of Page Sandstone; east of the depositional to erosional margin of the Page, the surface amalgamates with the J-2 and truncates the Navajo Sandstone. Locally, eolian sandstone of the Page is overlain by eolian deposits in the Upper Carmel Formation. There a crinkled surface or seam of bentonite marks the unconformity. West of the extent of the Page, the Upper Carmel rests directly on the Lower Carmel and the surface is overlain by discontinuous gypsum deposits. In SW Utah granule to pebble volcanic conglomerate lies on or just above the unconformity. Unconformity J-sup probably correlates with a sequence boundary recognized across the Northern Rocky Mountain region (Brenner and Peterson, 1994) and may mark the Bajocian–Bathonian boundary (Imlay, 1980). Regional stratigraphic and sedimentologic patterns suggest that the unconformity was caused by tectonic events, possibly related to events in the Cordilleran arc (Blakey et al., 1996). 4.5.11. Upper Carmel, Entrada The Upper Carmel Formation is a poorly studied but lithologically variable interval that was deposited across much of the Colorado Plateau (Figure 31). It contains red sandstone and mudstone of fluvial, sabkha, and eolian

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Figure 28 Upper Glen Canyon Group isopach and £uvial--eolian paleo£ow map (modi¢ed from Blakey et al., 1988). Note that £uvial--eolian facies boundary is nearly coincident with that of underlying sequence. The west-£owing Kayenta (sandy facies) and coeval Springdale £uvial system underlies the Navajo Sandstone north of the south margin of the Springdale Sandstone Member.

origin, marine limestone, fluvial volcanic-pebble conglomerate, and restricted marine gypsum (Blakey et al., 1996). The Entrada Sandstone overlies the Upper Carmel Formation. Although the contact has not been studied in detail, at many local outcrops it appears to be conformable and gradational. The Entrada Sandstone was deposited in widespread eolian erg complexes, adjacent to and inland from a restricted marine seaway (Peterson, 1994; Crabaugh and Kocurek, 1993). Although the Entrada is as widespread as the older Navajo eolian system, it does not approach the thickness of the Navajo (Figure 32). Regional facies patterns and stratigraphy of the Entrada are complex and reflect variations in depositional systems, subsidence rates, and height of the water table (Peterson, 1994; Crabaugh and Kocurek, 1993). The Upper Carmel Formation is Bathonian and the Entrada sandstone is Callovian (Imlay, 1980). The sequence thickens westward but is also truncated above by pre-Cretaceous erosion, so original maximum thickness is difficult to determine; isopachs shown on Figures 31 and 32 are minimums. The westward thickening was controlled by subsidence in the Utah–Idaho Trough, possibly an early phase of foreland basin development (Bjerrum and Dorsey, 1995). Abundant volcanic material, especially in the Upper Carmel Formation, reflects activity in the Cordilleran Arc (Blakey and Parnell, 1991; Chapman, 1989).

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Figure 29 Temple Cap isopach and facies map (modi¢ed from Blakey, 1994). The small remnant of this sequence is preserved between two regional unconformities.

4.5.12. Unconformity J-3 Unconformity J-3 forms a widespread erosional surface across much of the central and northern Colorado Plateau (Pipiringos and O’Sullivan, 1978). The unconformity is locally angular in central Utah, possibly due to soft-sediment deformation in the underlying Entrada Sandstone (Peterson, 1994). Like many of the Mesozoic unconformities of the Western Interior, it is best defined by regional patterns and is commonly difficult to locate on local outcrop, especially where strata underlying the surface are similar to strata above.

4.5.13. Curtis–Summerville The Curtis–Summerville interval is exposed across much of the central and northern Colorado Plateau (Figure 33); its extent is limited by erosion associated with the overlying J-5 and sub-Cretaceous unconformities (Peterson, 1988a). The sequence was deposited during the last Jurassic marine incursion into the Western Interior, and although sandstone and mudstone dominate the interval, carbonate content increases northward. The Curtis Formation is chiefly fine-grained, light-colored sandstone and sandy limestone and the overlying and partly coeval Summerville Formation comprises red, thin-bedded sandstone and mudstone. The former was deposited in a varied depositional setting of shallow marine and shoreline deposits and the latter was formed in

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Figure 30 Page-Lower Carmel isopach, facies, and eolian paleo£ow map (modi¢ed from Blakey et al., 1996). Note strong parallel trends between facies patterns and isopachs.

restricted marine and arid coastal plain settings (Peterson, 1994). The Curtis grades southeastward into eolian deposits of the Moab Tongue of the Entrada Sandstone and the Summerville grades southward into eolian and fluvial sandstone of the Romana Sandstone (Caputo and Pryor, 1991; Peterson, 1994). The Curtis–Summerville sequence marks the oldest interval of the study area that received sediment directly from uplifted Paleozoic rocks of the Cordilleran region (Peterson, 1988b), further evidence of uplift and thrusting to the west during the Jurassic (Bjerrum and Dorsey, 1995). Thus, from Early to Late Jurassic, stream deposits reflect a 1801 shift of stream flow and change in source from Pangean terranes at the beginning of the Jurassic to Cordilleran terranes at the end of the period (Blakey, 1994).

4.5.14. Unconformity J-5 The J-5 unconformity is another prominent and significant Early Mesozoic sequence boundary, perhaps second only to the J-2 (Pipiringos and O’Sullivan, 1978). The paleogeology below the surface suggests tectonic warping of the western craton. Currie (1998) suggested that the surface was formed by a migrating forebulge system that signaled early phases of thrusting in the Sevier orogenic belt. The J-5 is best understood through regional study as at many places strata under the surface are nearly identical to strata above (Peterson, 1988a). Controversy surrounds the unconformity, especially in the Four Corners region. Regional stratigraphic studies undertaken by

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Figure 31 Upper Carmel isopach, facies, and £uvial--eolian paleo£ow map (modi¢ed from Blakey et al., 1996). Note strong shift in £uvial paleocurrents relative to older Jurassic and Triassic rocks.

the U.S. Geological Survey (see summary in Peterson, 1988a) demonstrated that significant eolian deposits occur on either side of the surface; post-J-5 eolian deposits in the Morrison Formation represent erg sequences formed down wind from semiarid Morrison streams. The Bluff Sandstone, Junction Creek Sandstone, and Recapture Member of the Morrison Formation contain the largest of these eolian deposits. In contrast, Anderson and Lucas (1994) argued that no significant eolian deposits occur above the J-5 and that rather the above units are part of the Entrada and Summerville intervals and the J-5 occurs higher in the section. My own experience with these rocks strongly supports the U.S. Geological Survey correlations; eolian deposition in the Western Interior ended during the deposition of the Morrison, not before.

4.5.15. Morrison formation and younger Mesozoic events The J-5 unconformity marks the top of the interval of study for this chapter; however, a brief summary of subsequent Mesozoic events is presented to place J-5 and older events in a broader context. The Upper Jurassic Morrison Formation clearly defines the beginning of the dominance of westerly derived coarse-grained detritus onto the Colorado Plateau–Southern Rocky Mountain region. The Morrison has been interpreted as a foreland basin deposit to backbulge deposit (DeCelles and Currie, 1996) associated with Sevier thrusting in the Cordilleran region. Significant uplift of the SW margin of the Colorado Plateau followed Morrison continental

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Figure 32 Entrada (pre- J-3) isopach, facies, and eolian paleo£ow map (modi¢ed from Blakey et al., 1988). Note dramatic westward thickening adjacent to westward truncation by Cretaceous rocks.

deposition to form the pre-Cretaceous (K-0) unconformity; the erosional surface that resulted cuts down section dramatically from south-central Utah to SW Utah and removed perhaps 1,000 m of older Jurassic deposits (Peterson, 1988a; Blakey, 1989). The complex subsequent Cretaceous deposition filled the rapidly subsiding Rocky Mountain foreland basin with 1,000s of m of siliciclastic Cretaceous deposits that were derived from the Sevier Orogenic Belt (Miall et al, Chapter 9, this volume). The Rocky Mountain foreland basin area was then subjected to Laramide tectonics during latest Cretaceous and Early Tertiary with sharp uplifted blocks and adjacent foreland basins. Given that the youngest Cretaceous deposits formed in fluvial systems that graded to the retreating Western Interior seaway, all of the several kilometers of uplift that affected the Colorado Plateau and Southern Rocky Mountains took place in latest Cretaceous and Cenozoic.

5. Tectonic Origins of Pennsylvanian–Permian Basins 5.1. Introduction Pennsylvanian and Permian basins of the southern Western Interior are difficult to classify by modern classification systems. Most modern basin classification is based partly or entirely on origin and geologic setting

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Figure 33 Curtis--Summerville isopach, facies, and £uvial--eolian paleo£ow map (modi¢ed from Peterson, 1988a). East£owing streams sourced the mountains of the coeval Nevadan orogeny.

(Busby and Ingersoll, 1995) and because the basins of the Ancestral Rockies have controversial origin and setting, in relation to Late Paleozoic tectonics (e.g. Ye et al., 1996), they do not readily fit into modern classification. Therefore, I will use a simple scheme to discuss and classify the basins of the Ancestral Rockies region. I divide them into three general types, (1) yoked, (2) non-yoked, and (3) Cordilleran. The general structure and tectonics and sedimentary history of each is summarized below.

5.2. Yoked basins Yoked basins (Kay, 1951) are immediately adjacent to sharp Ancestral Rocky Mountain cratonic uplifts (Uncompahgre, San Luis, and Front Range uplifts) and their subsidence history and basin fill are closely related (yoked) to the uplift history of the positive elements (Kluth, 1986; Sloss, 1988). These basins display one or more, sharp, fault-bounded margins and coarse-grained detritus derived from the adjacent uplift. Structural relief with adjacent uplift is measured in thousands of meters and preserved Late Paleozoic deposits approach or exceed 3,000 m (Figure 5). Yoked basins within the study area are the Paradox Basin, Central Colorado Trough–Eagle Basin, Taos Trough, and ancestral Denver Basin. Yoked basins parallel paleotopographic and structural trends of the Uncompahgre, San Luis, and Front Range uplifts, generally elongated NW-NNW and at right angles to the

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Ouachita–Marathon fold and thrust belt. Hoy and Ridgway (2002) reviewed previous studies and presented new evidence that documented that some Ancestral Rockies faults are thrust faults (see their Figure 1) and that the Central Colorado Trough and Paradox Basin are flexural foreland basins and that parts of the Aphishapa and Sierra Grande uplifts are forebulges. They did not speculate as to plate tectonic style or setting for the Ancestral Rockies. Other workers have speculated on the tectonic setting and causes of the greater Ancestral Rocky Mountains system. In general, these speculations fall into three hypotheses. The first (Kluth and Coney, 1981; Kluth, 1986; Dickinson and Lawton, 2003) argued for a close cause and effect relationship between the Ouachita–Marathon and Ancestral Rocky Mountains systems. These workers cited similarities between timing, rates, and distribution of tectonic events, such as the east to west and outboard to inboard migration of Ancestral Rockies events that parallel similar patterns in the Ouachita–Marathon system. To further their arguments, they cited the ‘‘unique’’ setting of the Ancestral Rockies; they lay on the SW peninsular projection of SW North America and the Transcontinental arch. Such a precarious position coupled with tectonic events in the Ouachita–Marathon system may have forced SW North America northward relative to the continent as a whole to produce the uplifts and related (yoked) basins of the Ancestral Rockies system (Kluth and Coney, 1981). A second hypothesis, (Stevenson and Baars, 1986) cited structural orientation and style, and sedimentary sequences as evidence that the Ancestral Rockies system was part of a continent-wide wrench-fault system, probably related to complex Late Paleozoic tectonic events across southern North America. Some of the basins were cited as large-scale examples of pull-apart basins. The third hypothesis (Ye et al., 1996), discounted a close tectonic cause and effect between the Ouachita– Marathon system (or the more easterly Appalachian system) and the Ancestral Rockies. They also argued that the tectonic style of the Ouachita–Marathon system, a Carpathian-style orogeny rather than continental collision, was unlikely to trigger tectonic events of the Ancestral Rocky Mountains. Carpathian-style orogenies feature a fast-moving upper plate that comprises oceanic and arc crust or thin continental crust and rarely involve extensive metamorphism or basement-involved uplift. To account for the Ancestral Rocky Mountains, they invoked a relationship between a poorly known Andean-style arc that affected northern Mexico during the Desmoinesian and Permian as the triggering mechanism for Ancestral Rocky Mountain deformation (see Figure 2B). They compared the tectonic causes and settings of the Ancestral Rockies to those of the Laramide Rocky Mountains, involving shallow subduction of vast plates of the paleo-Pacific region beneath western North America. The above works and references cited therein demonstrate the wide-ranging opinions as to structural style and tectonic settings of the Ancestral Rockies system. Perhaps the greatest cause for this disparity is that key elements of the Ancestral Rockies were rejuvenated by later Laramide tectonics (see discussion in Hoy and Ridgway, 2002). The sedimentary fill of yoked Ancestral Rockies basins is marked by rapidly deposited sediment and contrasting lateral and vertical facies changes closely tied to adjacent tectonic events of the adjacent yoked uplifts (Hoy and Ridgway, 2002 and references cited therein). Subsidence patterns were complex, locally very rapid, and not completely synchronous from basin to basin; Kluth and Coney (1981) suggested a general progression of time-transgressive events from SE to NW across the region. In general, grain size and siliciclastic content of sedimentary facies reflect proximity and relief of adjacent uplifts. A general model of yoked-basin deposition is presented in Figure 34.

5.3. Non-yoked basins Non-yoked basins comprise an important component of the greater Ancestral Rocky Mountains system and they contrast with yoked basins in several important ways. Coarse-grained arkosic sediment is reduced or absent, total preserved thickness of Pennsylvanian and Permian deposits is thinner, subsidence rates were slower, basins lack major faults and folds at their margins, and lie adjacent to arches or low uplifts rather than major uplifts (Figure 5). Within the study area, the non-yoked basins are the Sweetwater Trough, Orogrande Basin, Pedregosa Basin, Holbrook Basin, and Grand Canyon Embayment. Generally excellent outcrop and/or subsurface data have yielded many detailed stratigraphic studies of these basins, but unfortunately, the subtle Late Paleozoic tectonics within the basins make interpretations of mechanisms of basin subsidence difficult to determine. Transpressional, transtensional, and foreland mechanisms have each been invoked (Soreghan, 1994). Sedimentation in non-yoked basins (Figure 35) reflects the general tectonic setting, especially with regard to basin edge, shelf, slope, or center (Ross, 1973; Blakey and Knepp, 1989). Soreghan (1994) followed models proposed by Heckel (1991) to demonstrate that much of the cyclicity within the Pedregosa and Orogrande basins was strongly controlled by glacioeustacy. Blakey and Middleton (1983) reached similar conclusions for cyclic eolian and sabkha seposits in the Holbrook Basin. Blakey and Knepp (1989) and Blakey (1990) demonstrated the non-synchronous nature of subsidence between basins in Arizona and that substantial subsidence occurred in

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Figure 34 Block stratigraphic diagram showing facies, paleotectonic, and paleogeographic setting, and depositional systems in a typical Pennsylvanian yoked basin. The basin shown is the southern portion of the Taos trough during the Desmoinesian; based on Casey (1980). The bounding fault may have a thrust component (see Hoy and Ridgway, 2002).

Figure 35 Block stratigraphic diagram showing facies, paleotectonic, and paleogeographic setting, and depositional systems in a typical Permian non-yoked basin. The basin shown is the Grand Canyon embayment during the Wolfcampian; from Blakey (1990, 1996).

some basins well after the generally accepted end of Ancestral Rockies orogeny. For example, the Grand Canyon Embayment underwent maximum rates of subsidence in the Wolfcampian (near the end of Ancestral Rockies uplift); however, the Holbrook Basin only underwent significant subsidence during the Leonardian — no evidence of basin activity was present before or since that time (Blakey, 1990). The Pedregosa Basin received thick Desmoinesian, Missourian, Wolfcampian and Leonardian sedimentation (Armin, 1987; Blakey and Knepp, 1989). The non-yoked basins are separated from each other by structurally and topographically higher regions generally referred to as arches, uplifts, or upwarps (Kluth and Coney, 1981; Armin, 1987; Blakey and Knepp, 1989), referred to as arches throughout remainder of this chapter. Pennsylvanian deposits are thin to absent across

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the arches, commonly pinching out across the structures; Permian deposits are present across each and reflect varying rates of subsidence across the features (Figure 4). Sedimentary facies patterns in regions dominated by non-yoked basins and arches are complex and vary rapidly both vertically and laterally (Figures 4–14). Carbonate and siliciclastic facies and marine and non-marine environments are present in all non-yoked basins. In some cases thin marine marker beds, usually carbonates, extend long distances from basins onto arches and provide key marker beds in otherwise generally nonfossiliferous deposits (Figure 4). Eolian dune environments are particularly well represented in regions south and west of the Uncompahgre Uplift and facies patterns associated with these sandstone units are particularly complex (Blakey et al., 1988; Blakey, 1990; Blakey, 1996).

5.4. Cordilleran basins Cordilleran basins lie to the west of the Ancestral Rocky Mountains but share strong relations to them and are briefly discussed here. Bissell (1970, 1974) provided stratigraphic syntheses of Pennsylvanian and Permian deposits of the eastern Cordilleran region and emphasized the great thickness and extent of Late Paleozoic sedimentation. Jordan and Douglass (1980) and Geslin, 1998 emphasized tectonic control on the enormously thick sedimentary accumulation; the complexity of these basins and their sedimentary fill is beyond the scope of this chapter and interested readers are referred to the above references.

6. Tectonic Setting of Triassic Basins 6.1. Introduction Triassic sedimentary rocks are widespread across the southern Western Interior. Although sedimentary facies are locally similar, overall stratigraphy and tectonic setting contrast sharply with underlying Upper Paleozoic rocks. The Ancestral Rocky Mountains were greatly reduced in relief and area and had relatively minor affect on sedimentary patterns except immediately adjacent to persisting topographic highs (Dubiel, 1994). Many Triassic basins show little relationship to older Ancestral Rockies elements; compare Figures 5, 17, and 18. Across much of the study area, Lower and Upper Triassic rock mirror each other in thickness trends as the former thicken to the NW and the latter thicken to the SE (Figure 16B).

6.2. Moenkopi shelf Lower and Middle Triassic sedimentation across most of the Colorado Plateau formed on a broad, extensive shelf that extended westward from the remnants of the Ancestral Rockies to the Wasatch line (Figure 17). Stewart et al. (1972b), Blakey (1974), Dubiel (1994), and Blakey et al. (1993) have detailed sedimentary patterns and regional stratigraphy of the Moenkopi Formation. Sedimentation began in the Lower Sythian and continued through several depositional sequences into the Middle Triassic (Blakey et al., 1993). At least as viewed from the Colorado Plateau, sedimentation took place in an extremely stable and simple tectonic setting, a tectonic lull between more complex settings.

6.3. Eastern Cordilleran Basin West of the Colorado Plateau, Lower and Middle Triassic deposits thicken dramatically across the Wasatch line (Paull and Paull, 1993; Blakey et al., 1993) (Figure 17). Saleeby and Busby-Spera (1992), related the westward thickening to variable tectonic events in the Cordilleran region including the Sonoman orogeny, the Golconda thrust, and Early Mesozoic development of Cordilleran subduction. See Ingersoll, Chapter 11 (this volume) for a thorough discussion of the Sonoman orogeny. The eastern Cordilleran basin is dominated by thick carbonate and dark mudstone deposits (Paull and Paull, 1993; Dubiel, 1994). Several carbonate tongues extend eastward onto the Colorado Plateau as members of the Moenkopi Formation and probably document marine highstands (Blakey et al., 1993).

6.4. Pre-Shinarump paleovalleys and Shinarump deposits The regional unconformity between the Lower and Middle Triassic Moenkopi Formation and Late Triassic Chinle Formation is marked by several spectacular paleovalleys across the Colorado Plateau (Blakey and Gubitosa, 1983). The paleovalleys are of two general types and contain deposits generally not present outside of the

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Figure 36 Panel cross-sections of Shinarump paleovalleys showing sedimentary architecture of valley ¢ll. A, General panel showing relations between narrow and broad paleovalleys. B, Narrow paleovalley ¢lled primarily with coarse sandstone and conglomerate based on outcrops inVermilion Cli¡s, Arizona and Circle Cli¡s, Utah. C, Narrow paleovalley ¢lled with heterolithic sandstone, mudstone, and conglomerate based on outcrops in Monument Valley, Utah--Arizona, Circle Cli¡s, Utah, and White Canyon, Utah. The inclined dot pattern represents lateral accretion deposits; gray areas represent mudstone deposits. All panels compiled from outcrop photos and ¢eld sketches.

paleovalleys (Figure 21). Broad paleovalleys generally trend NW-N and are tens of kilometers wide and up to a few tens of meters deep. They expose most of the deposits of the conglomeratic very coarse-grained sandstone of the Shinarump Member of the Chinle Formation. Narrow paleovalleys either lay within the boundaries of broad paleovalleys or outside but adjacent to the broader features (Figure 36). Narrow paleovalleys display spectacular relief locally exceeding 75 m (Blakey, 1974). They are also filled with Shinarump deposits, although locally these deposits contrast with those in broader paleovalleys (Blakey and Gubitosa, 1983). Specifically, the deposits are more variable ranging from conglomerate to carbonaceous mudstone and display variable bedding and cut and fill geometry. The paleovalley cutting and later fill, primarily by braided stream deposits, are clearly two separate episodes between the Middle and Upper Triassic. The narrow paleovalley cutting likely represents a sharp lowering of base level followed by initial Chinle deposition and then a period of extreme valley widening and succeeding broad braided stream aggradation (Blakey and Gubitosa, 1983). Although tempting to relate these events to worldwide sea level lows, tectonic events in the Cordilleran region may have also been responsible for changing base level conditions.

6.5. Chinle Basin Upper Triassic sedimentation mainly took place in a broad sedimentary basin that covered much of the Western Interior (Dubiel, 1994). The basin had two general centers, one in NW Colorado and the other in East-central Arizona (Stewart et al., 1972a). Isopach trends suggest that the remnant Uncompahgre uplift acted as a backbone flanked by thicker deposits on either side (Figure 18). The Chinle Formation is overlain by the J-0 unconformity, which truncates the formation in a W-SW direction; this partly explains the western thinning, opposite to that of the underlying Moenkopi Formation. The Chinle Basin has been described as a back-arc basin by Blakey and Gubitosa (1983), Dubiel (1994), and Lawton (1994) but as indicated above, modified by post-Chinle tectonics. I suggest that the Chinle was once more extensive and thicker, especially along the western margin of the Colorado Plateau, and stream trends suggest that it probably flowed to the west of the Plateau where deposits are currently absent, except for an outlier near Currie, Nevada (Dubiel, 1994). Post-Chinle uplift related to back-arc

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processes probably produced the noticed westward thinning. Early Mesozoic uplift has been documented in SE California and adjacent Arizona (Reynolds et al., 1989; Walker et al., 1983), possibly related to Late Paleozoic– Early Mesozoic truncation of SW North America (Stone and Stevens, 1988). Chinle deposition and stratigraphy are extremely complex but all workers agree that deposition occurred in an exclusively continental basin. Streams flowed generally to the NW across the depocenter (Figures 21–23), but many variations in stream style and flow have been noted (see reviews in Blakey and Gubitosa, 1983 and Dubiel, 1994).

7. Tectonic Setting of Jurassic Basins 7.1. Introduction Jurassic sedimentary basins occupy a transition from older tectonic trends described above to foreland basin deposition that dominated Cretaceous sedimentary history (Miall et al., Chapter 9, this volume). In general, Jurassic rocks thicken abruptly westward along the present margin of the Colorado Plateau and are then truncated by sub-Cretaceous or sub-Tertiary unconformities (Figure 24). Thinning to the SW was more gradual and past extent of Jurassic rocks in that direction can only be estimated as the south-facing Jurassic escarpment is today well north of areas of previous deposition (Figures 25 and 26).

7.2. Zuni sag The Zuni sag is a subtle negative feature better reflected in facies patterns than in isopach trends (Figures 27, 28, and 32). The feature lies at the foot of the Mogollon slope, a controversial upland area (Bilodeau, 1986; Riggs and Blakey, 1993), and trends NW across the SW Colorado Plateau. This trend was the locus of NW-flowing Glen Canyon streams, that flowed into the Utah–Idaho trough (Blakey, 1994). The Zuni sag is also the site of extensive sabkha deposits in the Entrada Sandstone and a prominent fluvial–eolian facies change in the Morrison Formation (Blakey et al., 1988). The Zuni sag probably was related to back-arc subsidence to the Jurassic Cordilleran arc.

7.3. Utah–Idaho trough The dominant basin during Jurassic sedimentation across the Western Interior was the Utah–Idaho trough (Imlay, 1980; Peterson, 1994). The basin lies along and west of the Wasatch line. Jurassic rocks thicken sharply into the trough beginning during Glen Canyon deposition, but especially during deposition of the San Rafael Group (Bjerrum and Dorsey, 1995). The basin margin also marks prominent facies changes from dominantly eolian deposition to the east to dominantly sabkha-marine deposition to the west (Blakey et al., 1988; Peterson, 1994; Blakey, 1994; Blakey et al., 1996). Debate continues as to classification and origin of the Utah–Idaho trough. Bjerrum and Dorsey (1995) carried out a detailed numerical basin analysis to document early foreland basin development as the cause of subsidence. The subsidence was tied to foreland thrusting in early phases of the Nevadan orogeny to the west (Figure 2D). Others have argued that true foreland basin subsidence did not initiate until the Early Cretaceous and that the Utah–Idaho trough was related to dynamic back-arc subsidence (see discussion in Lawton, 1994).

8. Summary: Tectonic Evolution and Controls on Deposition 8.1. Tectonic sequence of events Southwestern North America underwent several changes in tectonic setting from the Early Paleozoic through the Mesozoic (Figure 2). The following events are summarized from Marzolf (1990), E. L. Miller et al. (1992), Burchfiel et al. (1992), and Saleeby and Busby-Spera (1992). From Late Precambrian through Devonian, the region was part of a passive margin formed after rifting between North America and Gondwana. The study area was located well within the confines of the North American craton and received chiefly shallow marine and shoreline sedimentation (Figure 2A). From the latest Devonian into the Mississippian an arc collided with the western margin of the continent causing the Antler orogeny and signifying a change in tectonic setting. During the Pennsylvanian, the Ancestral Rocky Mountains formed, although the tectonic setting or cause of this event is not clearly understood. The study area straddled and bordered Ancestral Rocky Mountains uplifts (Figure 2B).

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During the latest Paleozoic and earliest Mesozoic (Figure 2C), another arc collided with the west edge of the continent forming the Sonoman orogeny (Ingersoll, Chapter 11, this volume); events were well west of the study area. During the Jurassic, a series of a terranes collided with the western margin of the continent and the Cordilleran arc radically changed its configuration (Figure 2D). Collectively, these events produced the Nevadan orogeny, also well to the west of the study area. Sometime (or times) between the Pennsylvanian and Late Jurassic, the SW margin of North America was truncated and apparently transported to the SE into Mexico along the Sonoran–Mojave megashear; see Anderson et al. (2005) for multiple views concerning this controversial feature and the chronology of events associated with it. Continued terrane collision and arc reorganization generated the Sevier orogeny throughout the Cordilleran region and into the study area. The exact beginning of the Sevier orogeny, Middle Jurassic or Early Cretaceous, remains in debate. The entire western third of the United States was affected by Late Sevier and Laramide (Rocky Mountain) tectonic events from Late Cretaceous through the Eocene. West of the study area, Sevier and Laramide compressional tectonics and Late Cenozoic crustal extension that is still occurring in some regions, greatly modified the earlier tectonic events that are covered in this chapter. However, on the Colorado Plateau, later events had little effect on earlier events except for uplift and exposure of Late Paleozoic and Mesozoic rocks. In the Southern Rocky Mountains, Laramide structure has overprinted older Ancestral Rocky Mountain structures and has made interpretations of the latter extremely difficult. Devonian through Jurassic tectonic events marginal to the study area had various effects on its sedimentary history, especially in the behavior and flow of rivers. Dramatic sedimentological change appears in the Early and Middle Pennsylvanian, coincident with Ancestral Rocky Mountains tectonics and Late stages of the Antler and related cordilleran tectonic events; following two hundred million years of carbonate shelf sedimentation, siliciclastic sedimentation appeared across much of the Colorado Plateau and Southern Rocky Mountains region. Adjacent to the Ancestral Rocky Mountains, coarse arkosic deposits grade laterally into sandstone, mudstone, and carbonate; fluvial paleocurrents radiate from uplifted areas towards adjacent basins. By Late Pennsylvanian and Early Permian, sandstone and mudstone dominated all but the most distal parts of the region; fluvial paleocurrents suggest that rivers drained westward into the Cordilleran seaways and south and southeastward into basins of Texas, and southern New Mexico and Arizona. Carbonate sedimentation returned across the western portion of the study area during the last Permian marine highstands, but fine-grained redbeds and evaporates continued to persist across the eastern portion of the region. Triassic and Early Jurassic river systems of the study area reflect the first-order tectonic grain of North America: rivers flowed west and northwest from the Appalachian continental divide towards the paleo-Pacific Ocean. A sharp increase in grain size of externally sourced clasts is observed in the Upper Triassic Chinle Formation, reflecting changes in source-terrane tectonics, climate, base level, or perhaps all three. In fact, from a pure sedimentologic point of view across the Colorado Plateau, the big change occurred with the initiation of pre-Chinle paleovalleys and continued during subsequent Chinle deposition. Not coincidentally, this may mark the time when western North America underwent significant tectonic changes following the Sonoman orogeny (Ingersoll, Chapter 11, this volume). Westward thickening of Lower Triassic marine deposits possibly reflects back-arc or foreland basin tectonics related to the Sonoman orogeny. Back-arc basin development, dynamic subsidence, and possible early foreland basin subsidence are apparent across much of the Colorado Plateau during the Late Triassic and Jurassic (Lawton, 1994; Ingersoll, Chapter 11, this volume). The earliest swing in fluvial paleocurrents reflecting sources to the southwest and west is recorded in the Middle Jurassic Carmel Formation and Late Jurassic Summerville and Morrison formations (Blakey et al., 1996; Peterson, 1988a). These changes reflect the evolving and maturing Cordilleran arc and developing Nevadan orogeny to the west.

8.2. Climatic controls Coincident with evolving tectonic conditions across the study area were evolving climatic settings. The following is summarized from Blakey et al. (1988), Peterson (1984, 1988c), Parrish and Peterson (1988), Dubiel (1994), and Blakey (1994, 1996). During the Early and Middle Paleozoic, the region was generally within equatorial zones although specific climatic indicators such as evaporites, redbeds, and coals are absent. By Pennsylvanian time, most of the study area was within a large tropical desert on the western margins of Pangaea. Widespread eolian sandstone, redbeds, evaporites, and marine faunal elements confirm this setting. An arid setting continued throuout the Permian (redbeds, eolian sandstone, evaporites) and into the Triassic (redbeds, minor evaporites, minor eolian sandstone). However, by Late Triassic Chinle deposition, conditions had become more humid, probably monsoonal as documented by fluvial style and abundant flora. Redbeds and eolian sandstone in the Upper Chinle document a return towards more arid conditions. The entire Early and Middle Jurassic were dominated by arid conditions (widespread eolian deposits, redbeds, evaporites) that probably became more monsoonal in the Late Jurassic during Morrison deposition (flora, stream style, scattered eolian deposits). Humidity increased until by Middle Cretaceous, greenhouse conditions dominated the region.

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Figure 37 Chart comparing Pennsylvanian--Jurassic timescale (Gradstein et al., 2004) with sea level curves (Ross and Ross, 1988; Haq, 1991), Sloss sequences (Sloss, 1988) and the stratigraphic sequences and unconformities of this report.Vertical scale in millions of years. On the column, black represents unconformities and yellow depositional sequences. The beginning and ending of sequences are approximated from the multitudes of data used in preparing this report, and the sequences are, of course, full of lesser unconformities. Asterisks signify high-water periods on the Colorado Plateau as marked by marine sequences that penetrate the dominantly continental section: 1, Morrowan marine limestone, western Grand Canyon; 2, Atokan marine limestone, western Grand Canyon; 3, Desmoinesian marine carbonates, Fossil Creek, Arizona; 4,Virgilian marine limestone, western Grand Canyon; 5,Wolfcampian marine limestone (Pakoon Formation), western Grand Canyon and (Elephant Canyon Formation) east-central Utah; 6, Fort Apache Limestone Member, Schnebly Hill Formation, Mogollon Rim; 7, marine limestone, Toroweap Formation (Brady Canyon Member), western Colorado Plateau; 8, marine limestone Kaibab and San Andres formations, western and southern Colorado Plateau; 9,Timpoweap--Sinbad limestones;10,Virgin limestone, Moenkopi Formation, western Colorado Plateau; 11, 12, marine carbonates of Judd Hollow Member (penetrating eolian Page Sandstone), south-central Utah; 13, marine limestone marker, Upper Carmel Formation (Paria River Member), south-central Utah; 14, marine sandstone, Curtis Formation, central Utah. Note the lack of marine incursion during Sloss’s Absaroka 3 interval, a time span of almost 70 million years. In places across the southern Colorado Plateau, over 400 m of continental strata are preserved during this interval.This suggests that local--regional tectonic and sedimentologic conditions overrode the global eustatic signal.

8.3. Eustatic controls Eustatic controls played important roles on deposition of rocks of the study area and several marine marker beds that penetrate continental sediments document marine highstands. A summary of the stratigraphic record

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compared to global sea level charts of Ross and Ross (1988) and Haq (1991) suggests a general correlation (Figure 37). In the Pennsylvanian, Soreghan (1994) documented eustatic control on deposits in the Pedregosa and Orogrande basins and Goldhammer et al. (1994) have documented eustatic controls on sedimentation in the Paradox Basin. Blakey and Middleton (1983) documented eustatic controls on eolian-sabkha deposits of Permian age in the Holbrook Basin; the marine marker bed Fort Apache Member of the Schnebly Hill Formation may correspond to a Leonardian global highstand (Blakey, 1990). The moderately thick Triassic continental succession across the region corresponds to global sea level lows; Lupe and Silberling (1985) related sequences in the Chinle Formation to sea level cycles in western Nevada. The incredible thickness of Navajo Sandstone in the Utah– Idaho trough contains no known marine tongues, perhaps because the low Early Jurassic sea level could not flood even the rapidly subsiding continental basin. Middle Jurassic global highstands may be reflected in marine tongues in the Carmel and Curtis formations.

8.4. What was the dominant control? As the preceding discussion suggests, climate, eustasy, and tectonics all weighed heavily on Late Paleozoic and Mesozoic sedimentation of the Southern Rocky Mountains and Colorado Plateau regions. Tectonic setting and events orchestrated the source areas and sediment supply, formed the basins of accumulation, terminated depositional events, and controlled climatic setting. Eustasy shaped the sequences and their boundaries (Sloss, 1996), partitioned marine and non-marine environments, and provided accommodation space. Climate mediated and modified depositional systems, mitigated sediment supply, and produced climate-sensitive deposits, not the least of which was the production of the greatest preserved eolian accumulation in the history of the Earth. These and other subsidiary controls were all of paramount importance and to attempt to delegate control to a dominant factor is at present futile.

ACKNOWLEDGMENTS Although much of this chapter reflects my 35 years of work with the sedimemtary rocks on the Colorado Plateau, many of the data and conclusions are based on 35 Master’s theses and I am indebted to these students and their dedication, persistence, and scholarship. Discussions, debates, and collaboration with many other workers has greatly enriched my knowledge of the region; I especially thank Don Baars, Stan Beus, Margie Chan, Becky Dorsey, Lars Clemmensen, Bill Dickinson, Bob Dott, Russ Dubiel, Bill Furnish, Karen Havholm, Phil Heckel, Ray Ingersol, Chuck Kluth, Gary Kocurek, Bart Kowallis, Rip Langford, Tim Lawton, Dave Loope, Andrew Miall, Larry Middleton, Mike Morales, Bill Parker, Judy Parrish, Wes Peirce, Pete Peterson, Steve Reynolds, Nancy Riggs, Lynn Soreghan, Lee Stokes, Christine Turner, and Paul Umhoefer. Tim Lawton and Andrew Miall provided formal reviews of the manuscript.

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Poole, F. G., Stewart, J. H., Palmer, A. R., Sandberg, C. A., Madrid, R. J., Ross, R. J., Jr., Hintze, L. F., Miller, M. M., and Wricke, C. T., 1992, Latest Precambrian to latest Devonian time; development of a continental margin, in Burchfiel, B. C., Lipman, P. W., and

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Zoback, M. L. eds., The Cordilleran Orogen: conterminous U. S.: the geology of North America, Decade of North American Geology, Geological Society of America, Boulder, v. G-3, pp. 9–56. Rawson, R. R., and Turner-Peterson, C. E., 1980, Paleogeography of northern Arizona during the deposition of the Permian Toroweap Formation, in Fouch, T. D. and Magathan, R. eds., Paleozoic paleogeography of the west-central United States, Rocky Mountain Section, SEPM, pp. 341–352. Reynolds, S. J., Spencer, J. E., Asmerom, Y., DeWitt, E., and Laubach, S. E., 1989, Early Mesozoic uplift in west-central Arizona and southeastern California. Geology, v. 17, pp. 207–211. Riggs, N. R., and Blakey, R. C., 1993, Early and Middle Jurassic paleogeography and volcanology of Arizona and adjacent areas, in Dunne, G. and McDougall, K. eds., Mesozoic Paleogeography of the Western United States-II, Pacific Section SEPM (Book 71), pp. 347–375. Riggs, N. R., Lehman, T. M., Gehrels, G. E., and Dickinson, W. R., 1996, Detrital zircon link between headwaters and terminus of the Upper Triassic Chinle-Dockum paleoriver system. Science, v. 273, pp. 97–100. Ross, C. A., 1973, Pennsylvanian and Early Permian depositional history, southeastern Arizona. American Association of Petroleum Geologists Bulletin, v. 57, pp. 887–912. Ross, C. A., and Ross, J. R. P., 1988, Late Paleozoic transgressive–regressive deposition, in Wilgus, C. K., Hastings, B. S., Ross, C. A., Posamentier, H., and St. Kendall, C. G. C. eds., Sea-level changes: an integrated approach, SEPM (Special Publication 42), pp. 227–247. Saleeby, J. B., and Busby-Spera, C., 1992, Early Mesozoic tectonic evolution of the western U. S. Cordillera, in Burchfiel, B. C., Lipman, P. W., and Zoback, M. L. eds., The Cordilleran Orogen: conterminous U. S.: the geology of North America, Decade of North American Geology, Geological Society of America, Boulder, v. G-3, pp. 107–168. Silberling, N. J., and Roberts, R. J., 1962, Pre-Tertiary stratigraphy and structure of northwestern Nevada. Geological Society of America Special Paper 72, 58 pp. Sloss, L. L., 1988, Tectonic evolution of the craton in Phanerozoic time, in Sloss, L. L. ed., Sedimentary Cover — North American craton: U.S.: the geology of North America, Geological Society of America, Boulder, CO, v. D-2, pp. 25–53. Sloss, L. L., 1996, Sequence stratigraphy on the craton: caveat emptor, in Witzke, B. J., Ludvigson, G. A., and Day, J. eds., Paleozoic sequence stratigraphy: views from the North American craton, Geological Society of America (Special Paper 306), Boulder, CO, pp. 425–434. Soreghan, G. S., 1994, Stratigraphic response to geologic process: Late Pennsylvanian eustacy and tectonics in the Pedregosa and Orogrande basins, Ancestral Rocky Mountains. Geological Society of America Bulletin, v. 106, pp. 1195–1211. Stevenson, G. M., and Baars, D. L., 1986, The Paradox: a pull-apart basin of Pennsylvanian age, in Peterson, J. A. ed., Paleotectonics and sedimentation in the Rocky Mountain region, United States, American Association of Petroleum Geologists (Memoir 41), pp. 513–541. Stewart, J. H., Anderson, T. H., Haxel, G. B., Silver, L. T., and Wright, J. E., 1986, Late Triassic paleotopography of the southern Cordillera: the problem of a source area for the voluminous volcanic detritus in the Chinle Formation of the Colorado Plateau region. Geology, v. 14, pp. 567–570. Stewart, J. H., Poole, F. G., and Wilson, R. F., 1972a, Stratigraphy and origin of the Upper Triassic Chinle Formation and related Upper Triassic strata in the Colorado Plateau region. United States Geological Survey Professional Paper 690, 336 pp. Stewart, J. H., Poole, F. G., and Wilson, R. F., 1972b, Stratigraphy and origin of the Triassic Moenkopi Formation and related strata in the Colorado Plateau region. United States Geological Survey Professional Paper 691, 195 pp. Stone, P., and Stevens, C. H., 1988, Pennsylvanian and Early Permian paleogeography of east-central California: implications for the shape of the continental margin and the timing of continental truncation. Geology, v. 16, pp. 330–333. Trexler, J. H., Cashman, P. H., Jr., Snyder, W. S., and Davydov, V. I., 2004, Late Paleozoic tectonism in Nevada; timing, kinematics, and tectonic significance. Geological Society of America Bulletin, v. 116, pp. 525–538. Tweto, O., 1977, Tectonic history of west-central Colorado: in Exploration frontiers of the central and southern Rockies, Rocky Mountain Association of Geologists, Denver, pp. 11–22. Walker, J. D., Burchfiel, B. C., and Royden, L. H., 1983, Westward-derived conglomerates in the Moenkopi Formation of southeastern California, and their probable tectonic significance. American Association of Petroleum Geologists Bulletin, v. 67, pp. 320–322. Welch, J. E., 1958, Faunizones in the Pennsylvanian and Permian rocks of the Paradox basin, in Intermountain Association of Petroleum Geologists Guidebook, 9th Field Conference, Geology of the paradox basin, Salt Lake city, pp. 153–162. Wengard, S. A., and Matheny, M. L., 1958, Pennsylvanian System of the Four Corners region. American Association of Petroleum Geologists Bulletin, v. 42, pp. 2048–2106. Whalen, M. T., 1996, Facies architecture of the Permian Park City Formation, Utah and Wyoming: implications for the paleogeography and oceanic setting of western Pangea, in Longman, M. W. and Sonnenfeld, M. D. eds., Paleozoic systems of the Rocky Mountain Region, Rocky Mountain Section SEPM, pp. 353–378. Ye, H., Royden, L., Burchfiel, C., and Schuepbach, M., 1996, Late Paleozoic deformation of interior North America: the Greater Ancestral Rocky Mountains. American Association of Petroleum Geologists Bulletin, v. 80, pp. 1397–1432.

CHAPTER 8

The Southern Midcontinent, Permian Basin, and Ouachitas Andrew D. Miall

Contents 1. Introduction 2. Early to Middle Paleozoic Structural and Stratigraphic Setting 3. Cyclothems 4. Cyclic Sedimentation at the Shelf Margin 5. Ouachita Deformation and Sedimentation 6. The Permian Basin and the Capitan Reef 7. Evaporite Sedimentation in the Delaware and Midland Basins 8. Oil and Gas Production 9. Summary Acknowledgments References

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Abstract This chapter is concerned with the southern portion of the craton, the southern craton margin, and its geological relationship with the Ouachita Orogeny. The southern continental margin was shaped by extensional and transform faulting, during the breakup of Rodinia, and includes the Oklahoma Basin, a basin formed by transtensional faulting. Carbonate reef systems flourished along the entire continental margin from Texas to Newfoundland during the Early Paleozoic. Shallow-water, carbonate, and siliciclastic sedimentation continued in the Oklahoma Basin from Middle Ordovician through Earliest Mississippian time. The ocean that lay south of the North American Plate through much of the Paleozoic was closed by collision with Gondwana, starting in the Mid-Mississippian. Northwesterly directed contractional stress generated a regionally distinctive style of transpressive fault and block-uplift deformation across the Southwestern United States. Mild warping of the Texas–Oklahoma area during this period accentuated existing differentiation of the continent into a series of basins (including the Delaware and Midland basins) and intervening uplifts (Diablo Platform, Central Basin Platform, Eastern Shelf, Llano Uplift, Ozark Uplift). That portion of this basin and uplift system located within Texas now constitutes the classic Permian Basin. The carbonate platform and basin configuration of the Permian Basin is similar in topography and scale to that of the Cretaceous to modern Bahamas Platform. The basin is probably best known for the distinctive ‘‘clinoform’’ stratigraphy of the Capitan Reef, one of the first places where the lateral facies transition from back reef, through reef crest, foreef slope to basin floor was mapped and described in detail, based on superb outcrops in the Guadalupe Mountains. Sediments of the Absaroka Sequence (Uppermost Mississippian to Lower Jurassic) are widely distributed across the continental interior. For much of the Carboniferous to Permian, the Earth was under the influence of repeated, high frequency glacioeustatic sea-level changes. These caused transgressions and regressions that shifted shorelines and depositional systems back and forth hundreds to thousands of kilometers across low-relief continental interiors. The classic Pennsylvanian cyclothems of the southern US Midcontinent are the most characteristic and well-known products of this process. Crustal loading of the continental margin by colliding Gondwana terranes commencing in the Mid-Pennsylvanian generated typical foreland-type basins which now constitute the Ouachita system, and including such foreland basins as the Forth Worth, Arkoma, and Black Warrior basins. In the Ouachita Mountains of Oklahoma and Arkansas some 3,500 m of structurally disturbed Late Cambrian to Mississippian pre-orogenic strata have been mapped. They consist mainly of shales and sandstones, the latter representing a variety of sediment-gravity flow mechanisms.

Sedimentary Basins of the World, Volume 5 ISSN 1874-5997, DOI 10.1016/S1874-5997(08)00008-7

r 2008 Elsevier B.V. All rights reserved.

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1. Introduction The Midcontinent–Permian Basin area of the United States extends from North Dakota to Texas and from Colorado to Pennsylvania. Physiographically, this vast area includes most of the Mississippi–Missouri catchment, and represents, tectonically, most of the United States portion of the Interior Platform. The northern part of this area is underlain by the Great Plains; the southern part includes the plateau areas of Texas, Oklahoma, and Arkansas (e.g., Edwards Plateau, Ozark Plateau). The broad sedimentary history of this vast craton is described in Chapter 2. This chapter is concerned with the southern portion of the craton, the southern craton margin, and its geological relationship with the Ouachita Orogeny. Most of the details covered in this chapter deal with the first two subsequences of the Absaroka Sequence, which span the Late Mississippian to Late Permian (Sloss, 1963, 1988). Structurally, the southern margin of the Interior Platform is the Late Precambrian–Cambrian continental margin, which was formed by the breakup of the supercontinent Rodinia (Figure 1, Figure 6 in Chapter 1). The ancient divergent continental margin, now buried beneath younger rocks, extends northwestward across Alabama and Mississippi, and then bends southwestward across southern Texas and into Mexico. Extensional faulting accompanying Rodinia breakup during the Late Precambrian generated rift systems extending deep into the continental interior, including the Reelfoot Rift system beneath the Mississippi Valley, and the Oklahoma ‘‘aulacogen’’ — shown here in quotes, because current reconstructions do not support the original concept of this basin having developed at a triple-point junction (Brewer et al., 1983). The ocean that lay south of the North American Plate through much of the Paleozoic was closed by collision with Gondwana, starting in the Mid-Mississippian. Current reconstructions suggest that a series of amalgamated terranes, including a Yucatan Block, collided with the continental margin first (Figure 11 in Chapter 1), with full closure completed during the Permian. Northwesterly directed contractional stress generated a regionally

Figure 1 Structural elements of the southern mid-continent and Gulf Coast continental margin. Locations of structural and stratigraphic cross sections referred to in this chapter are shown by ¢gure number. Adapted from Cebull et al. (1976), Arbenz (1989) and Thomas (1989, 2006).

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distinctive style of transpressive fault and block-uplift deformation across the southwestern United States (an area corresponding approximately to the four-corners states), generically termed the Ancestral Rockies (Blakey, Chapter 7). Mild warping of the Texas–Oklahoma area during the same period accentuated existing differentiation of the continent into a series of basins (including the Delaware and Midland basins) and intervening uplifts (Diablo Platform, Central Basin Platform, Eastern Shelf, Llano Uplift, Ozark Uplift). That portion of this basin and uplift system located within Texas now constitutes the classic Permian Basin (Frenzel et al., 1988). Crustal loading of the continental margin commencing in the Mid-Pennsylvanian generated typical foreland-type basins which now constitute the Ouachita System, and including such foreland basins as the Forth Worth, Arkoma, and Black Warrior basins. Much of the Ouachita System is now covered by the Cretaceous– Tertiary cover of the Gulf Coast Province (Chapter 15). The Ouachita fold-thrust system is exposed in the Ouachita Mountains of Oklahoma and Arkansas and in the Marathon Uplift of Southwestern Texas, and the Anadarko Basin constitutes an important additional basin of foreland type, developed by contractional modification of an earlier rift basin, the Oklahoma ‘‘aulacogen.’’ Continental sedimentary patterns during Absaroka Sequence sedimentation are quite different from those of Earlier Paleozoic sequences (Sloss, 1988), displaying larger preserved sediment thicknesses and more rapid sedimentation rates in the southwestern quadrant of the continent than elsewhere, including the Ancestral Rocky Mountains area (Chapter 7) and the Permian Basin. This is partly due to the crustal loading processes associated with the Ouachita Orogen, but may also be due to dynamic topographic tilting of the continent. The Permian Basin and southern Midcontinent area is historically important as an area where many important geological concepts have been developed. Wanless and Weller (1932) coined the term cyclothem for the repetitive coal-bearing carbonate-clastic cycles of Pennsylvanian Age that had been observed throughout the mid-continent. Wanless and Shepard (1936) were the first to invoke repeated eustatic sea-level change as a mechanism for cyclic sedimentation, suggesting that the cyclothems were the product of high-frequency glacioeustastic transgressions and regressions that occurred as a result of a major (Devonian–Permian) continental glaciation centered on Gondwana. The same cyclic sea-level changes at the shelf margin generated a distinctive architecture of progradation, aggradation, and retrogradation in the Pennsylvanian–Permian section of Central Texas, which formed the basis for one of the earliest formal models of continental-margin cyclic sedimentation. The ‘‘depositional topography’’ concepts of Van Siclen (1958) were amongst the most important precursors of modern sequence stratigraphy. A study of Upper Pennsylvanian strata in New Mexico by Wilson (1967) developed Van Siclen’s ideas further, and introduced the concept of ‘‘reciprocal sedimentation,’’ a term which he proposed for the pattern of highstand carbonate and lowstand clastic sedimentation so characteristic of many mixed clastic–carbonate shelf margins. The carbonate platform and basin configuration of the Permian Basin is similar in topography and scale to that of the Cretaceous to modern Bahamas Platform (Scholle, 2006). The basin is probably best known for the distinctive ‘‘clinoform’’ stratigraphy of the Capitan Reef, one of the first places where the lateral facies transition from back reef, through reef crest, foreef slope to basin floor was mapped and described in detail, based on superb outcrops in the Guadalupe Mountains (King, 1948). The undaform–clinoform–fondoform concept of Rich (1951) was not based on any studies of the Capitan Reef, but since the work of King (1948) this has become one of the most frequently cited examples of clinoform architecture. One of the first, and much-cited, detailed studies of an ancient delta was carried out on the Bartlesville sandstone of Kansas and Oklahoma by Visher (1968) and Visher et al. (1971) (although interpretations have since changed, as discussed later). This unit has yielded petroleum since the first productive well was drilled in 1893, and has provided many useful studies of fluvial stratigraphic traps (e.g., Berg, 1986) and traps formed in incised valleys (e.g., Lyons and Dobrin, 1972; Ye and Kerr, 2000). The area has received much attention from geologists because of its petroleum productivity. Several major oil fields in the Bartlesville Sandstone were discovered in Oklahoma between 1897 and 1927. Extensive exploration of the Permian Basin began in the 1930s. Scholle (2006) reported that the basin is one of the most prolific in North America, containing some 91.6 billion barrels of oil in place, and 3 trillion cubic feet of natural gas. Production has come from many stratigraphic levels, ranging from the Cambrian to the Cretaceous, but the Permian Section contains some 71% of the total discoveries. Seven of the 20 largest oil fields in the continental United States are located in the Permian Basin. Up to the late 1980s some 20% of US hydrocarbon production was from the Permian Basin (Frenzel et al., 1988).

2. Early to Middle Paleozoic Structural and Stratigraphic Setting Reconstruction of the Eastern North American continental margin suggests that the portion of the margin crossing Texas from southwest to northeast is a typical rifted margin, while that portion which extends

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southeast across Mississippi to Florida is a transform margin (Figure 1, Figure 6 in Chapter 1; Cebull et al., 1976). Poole et al. (2005) documented a comparable rift-transform margin extending westward into Northern Mexico. From Cambrian or Late Precambrian time until the Early Mississippian, sedimentary and tectonic events on the continent north of this margin were typically cratonic in character. A broad sag developed over the southern continent beneath the area now corresponding approximately to the Delaware and Midland basins (Figure 2). This area, referred to as the Tobosa Basin, contains up to at least 1,200 m of Upper Cambrian to Late Devonian strata (Phanerozoic sedimentation along the continental margin commenced in the Late Cambrian). The deposits consist mainly of shallow-water carbonates and sandstones (Frenzel et al., 1988). The Tobosa Basin was essentially filled with sediment by the end of the Devonian, and then ceased to exist as a separate structural entity (Frenzel et al., 1988). Epeirogenic movements and in-plane stress related to the Ouachitan Orogeny, described later, subsequently modified the region, generating the tectonic configuration of the Permian Basin that persists to the present day. The Oklahoma Basin (Figures 1, 2) was initiated during continental breakup late in the Precambrian. This early basin, the depocentre of which is located beneath the present Texas–Oklahoma border, has been called an aulacogen, but modern plate-tectonic reconstructions (Figure 1) suggest that it is aligned with a fracture zone and, like the modern Benue Trough of Nigeria (also termed an aulacogen in early plate-tectonic literature), probably originated as a transtensional rift. Unusually for a basin of rift origin, the early sedimentary fill, a W2 km-thick succession of Middle Cambrian and Early Ordovician beds (Arbuckle Group), consists largely of limestone. These beds thin out to zero over an uplift in Northeast Oklahoma, which marked a positive area within what was otherwise a vast cratonic shelf extending across most of the continental interior at this time. Carbonate reef systems flourished along the entire continental margin from Texas to Newfoundland during the Early Paleozoic. Shallow-water, carbonate and siliciclastic sedimentation continued in the Oklahoma Basin from Middle Ordovician through Earliest Mississippian time, although subsidence appears to have been slow, and the section is interrupted by numerous unconformities. A distinctive feature of the continental margin is the Llano Uplift, an exposure of Precambrian igneous and metamorphic rocks that physiographically defines the southern margin of the Eastern Shelf, in Central Texas. The uplift is a structural dome that was stabilized by massive granite intrusions in late Precambrian time (Walker, 1992). It was intermittently active throughout the Paleozoic. The effect of the Ouachita Orogeny, during the Early and Middle Pennsylvanian, was to fragment the uplift into a series of horsts and grabens striking parallel to the leading edge of the thrust belt.

Figure 2 Isopach of Middle Ordovician sediments. This interval was selected to illustrate the typical pattern of continental margin sedimentation during the Early Paleozoic. Adapted from Cook and Bally (1975).

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In the Late Devonian much of the southern continent was affected by a transgression in which a dark, highly organic, fossiliferous shale was developed. In the Permian Basin this is assigned to the Woodford Formation, which reaches 120 m in thickness in the center of the Tobosa Basin. This unit correlates with the well-known Chattanooga Shale of Tennessee. Slope and basin sediments deposited along the southern margin of the continent are now exposed in the Marathon orogenic belt of Southwest Texas (Figure 3; Ross, 1986; Wuellner et al., 1986) and in the Ouchita Mountains (Lowe, 1989). They range in age from Cambrian to Mississippian, and consist of sandstones, siltstones, shales, calcarenites, limestone–pebble conglomerates, and boulder olistostromes. The siliciclastics and detrital carbonates were deposited primarily by sediment-gravity flows derived from the continental margin to the north. Dark shales contain a typical oceanic assemblage of graptolites, sponge spicules, and conodonts. In the Ouachita Mountains some 3,500 m of structurally disturbed Late Cambrian to Mississippian preorogenic strata have been mapped (Lowe, 1989). They consist mainly of shales and sandstones, the latter representing a variety of sediment-gravity flow mechanisms. In the easternmost portion of the exposed orogen, in North Central Arkansas, facies trends indicate the probable existence of a major submarine canyon feeding detritus into the ocean. Paleocurrent data indicate a southwesterly orientation for this channel, with paleoflow elsewhere in the mountains confirming a generally southwest to westerly transport direction for the turbidity currents and other sediment-gravity flows. The sandstones vary from quartzofeldspathic to lithic in character, and contain paleocurrent indicators, such as flute marks, consistent with derivation from the craton and passive continental margin to the north and northeast. A chert-rich unit, the Arkansas Novaculite, of probable Late Devonian to Early Mississippian Age, occurs at the top of the pre-orogenic succession in the Ouachita Mountains. Sponge spicules and radiolarians are locally abundant, suggesting slow, deep-water sedimentation far from detrital sediments sources. However, some of the beds preserve primary hydrodynamic structures and appear to have developed by replacement of currentdeposited sandstones and turbidites. In the Marathon Uplift, an Upper Mississippian turbidite sandstone unit (Tesnus Formation) shows a significant increase in grain size relative to earlier units, and contains abundant metamorphic and granitic rock fragments. Paleocurrent directions indicate sources to the southeast. This unit is interpreted as having been derived from tectonically active sources and is regarded as the first indicator of the approaching terranes that would generate the Ouachitan Orogeny (Figure 11, Chapter 1). The Lower and Middle Pennsylvanian Dimple Limestone formation occurs within an allochthonous thrust sheet in the Marathon Uplift. It consists of cherty limestone, shales, conglomerates, and bedded chert. Shelf, slope, and basinal facies are all represented. The unit is interpreted as the deposit of a carbonate atoll within the remnant Rheic ocean basin.

Figure 3

Structural elements in Southwest Texas (Ross, 1986).

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3. Cyclothems Sediments of the Absaroka Sequence (Uppermost Mississippian to Lower Jurassic) are widely distributed across the continental interior. Figure 4 provides a generalized facies map of Upper Pennsylvanian strata, but is not a particularly useful reflection of regional stratigraphic controls or the evolution of the paleogeography because, for much of the Carboniferous to Permian, the Earth was under the influence of repeated, highfrequency glacioeustatic sea-level changes (Crowell, 1978). These caused transgressions and regressions that shifted shorelines and depositional systems back and forth hundreds to thousands of kilometers across low-relief continental interiors. The Pennsylvanian cyclothems of the Southern US Midcontinent are the most characteristic and well-known products of this process. The original Midcontinent cyclothems have been the focus of numerous studies of cyclic sedimentation and of climatic and eustatic controls on sedimentation, since the glacioeustasy interpretation was first proposed by Wanless and Shepard (1936), and have provided the basis for interpretations of numerous other cyclic successions, including contemporaneous coal-bearing successions in Western Europe (e.g., Ramsbottom, 1979; Ross and Ross, 1988). The classic cyclothem model is illustrated in Figure 5, with an interpretation provided by Moore (1964). A convenient point to use for the definition of individual cyclothems is the disconformity surface that characteristically occurs at the base of a nonmarine succession of sandstone, shale, and coal. Incision by fluvial erosion is common at these disconformity surfaces. The nonmarine section typically displays an upward decrease in grain size and an upward increase in marine influence, culminating in a fully marine shale–limestone interval. The top of the cyclothem consists of a prograding deltaic complex, although the coarse, fully nonmarine top of the deltaic succession, and part of the underlying open-marine section, may be removed by erosion that forms the succeeding disconformity. The succession and interpretation shown in Figure 5 relate to the classic cyclothems of Kansas and Oklahoma. Elsewhere within the Midcontinent the succession varies, those occurring in the eastern part of the continental interior containing significantly greater clastic intervals, whereas those further west are dominated by carbonates. Heckel (1980, 1994) situated these regional variations within a continent-wide paleogeographic context (Figures 6, 7). The so-called ‘‘Appalachian cyclothems’’ are almost exclusively clastic in composition, and were clearly dominated by the sediment supply and the tectonic influence of the active Alleghanian Orogeny (Klein and Willard, 1989; Ettensohn, Chapter 4). The ‘‘Illinois cyclothems’’ are those that characterize the Illinois Basin, and consist of a mixed carbonate–clastic succession. Kansas cyclothems may be bundled into megacyclothems (Figure 6) indicating the simultaneous occurrence of cyclic sea-level changes at different frequencies. At the edge of the continent, in the deepest-water cratonic basins (Midland, Delaware basins), cyclothems consist largely of

Figure 4 Upper Pennsylvanian (Missourian) lithofacies (Cook and Bally, 1975).

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Figure 5

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The standard cyclothem model, with an interpretation suggested by Moore (1964).

shales. Along the western margin of the Permian Basin shelf, cyclothems are again clastic-rich, reflecting sediment sourcing from uplifts of the Ancestral Rocky Mountains. Synthesis of a very large amount of biostratigraphic and facies data provided Heckel (1980, 1986, 1994) with the basis for a detailed paleogeographic (Figure 7) and chronostratigraphic (Figure 8) interpretation for these rocks. During deep-water phases, a dark, conodont-bearing shale was deposited over wide areas of the

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Figure 6 Schematic cross-section through a ‘‘megacyclothem’’ showing facies variations from southwest to northeast across the continental interior (Heckel, 1994).

Figure 7 Paleogeography of the Midcontinent during a phase of maximum marine transgression during the Late Pennsylvanian (Heckel, 1994).

midcontinent, extending as far north as Nebraska, Iowa, and Illinois. During regressions, deltaic complexes prograded westward from the Appalachians along a belt extending from Pennsylvania to Oklahoma while, in the continental interior, conditions for the deposition of shelf limestones became widespread. More than 50 individual cycles have been correlated across the continental interior (Figure 8). These cycles were deposited over a time span of about 9 Myr, according to current age designations for the Desmoinesian, Missourian, and Virgillian (www.stratigraphy.org), indicating that they each represent an average of about 170,000 years. Heckel (1986) subdivided the cycles into major and minor varieties, depending on their regional extent and interpreted

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Figure 8 Sea-level curve for part of the Middle--Upper Pennsylvanian cyclothem sequence of the North American midcontinent, which ranges from 260 m thick in Iowa to 550 m thick in Kansas.The lateral extent of erosion surfaces at the base of the cycles is shown by diagonal hatching. Fluvial deltaic complexes are shown by a series of dots, and conodont-bearing shales by lines with dots (Heckel, 1986).

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time span, and suggested correlations with Milankovitch periodicities, but such correlations must remain speculative until more is known about the astronomical history of Earth that far back in time. However, the interpretation of the cyclothems as the product of orbitally forced glacioeustasy is by now the standard model.

4. Cyclic Sedimentation at the Shelf Margin Surface and subsurface exploration work during the 1940s and 1950s distinguished shelf, slope, and cratonic basin sediments and environments along the southern margin of the Midcontinent area, work that led to the definition of the major tectonic elements in the area, including the Midland and Delaware basins and the platform and uplift areas surrounding them (e.g., King, 1948). These tectonic elements are shown in Figure 1, and stratigraphic columns for the major basins and platform areas are shown in Figure 9. A landmark study by Van Siclen (1958) made explicit reference to Rich’s (1951) three ‘‘critical environments of deposition’’ in the Permian stratigraphy of Central Texas, at the edge of the Eastern Shelf, and developed a model for sedimentation under conditions of cyclic sea-level change that anticipated the Exxon sequence models by 30 years. Van Siclen (1958) showed that shelf-margin (‘‘undaform’’) shelf limestone units a few tens of meters thick could be traced through shelf-margin reefs and then on to depositional (‘‘clinoform’’) slopes of 1.5–51, down which they changed facies into siliceous shales, and then into nearly horizontal, black shales (petrophysically highly resistive) of basinal (‘‘fondoform’’) environments (Figure 10). In some cases, as in that shown in Figure 10, the positions of the changes in slope across the shelf margin shift gradually basinward through time, but unit thicknesses remain fairly constant from shelf to basin, indicating a combination of progradation and aggradation. In other cases, progradation is dominant, individual slope limestone units wedging out at the top of the clinoform. Van Siclen’s sedimentary model (Figure 11) showed how varying patterns of sea-level change combined with variations in sediment supply could generate all the observed variations in basin–margin architecture, and he cited the work of Wanless and Weller (1932) and Wanless and Shepard (1936) in suggesting repeated eustatic sea-level change as the cause of this cyclicity. Many subsequent studies of Permian Basin stratigraphy have emphasized the cyclicity of shelf and basin successions (e.g., Silver and Todd, 1969; Galloway and Brown, 1973), indicating the unusual importance of high-frequency sea-level changes during the Late Paleozoic. Uplift of the Wichita–Amarillo fold-thrust belt and the Ouachita Orogen during the Late Pennsylvanian led to erosion and the shedding of detritus south and west into the Eastern Shelf, at the margins of the Midland Basin. A study by Galloway and Brown (1973), which examined the Upper Pennsylvanian–Lower Permian succession at the shelf margin differentiated three distinct but interfingering depositional systems, the Cisco fluvial–deltaic system on the shelf, the Sylvester shelf-edge carbonate bank system, and the Sweetwater clastic slope system (Figure 12). Fluvial and deltaic deposits show similarities to those of the modern Mississippi system, and subenvironments comparable to those of the Mississippi were identified by Galloway and Brown (1973), including delta-plain, distributary mouth-bar, crevasse-splay, and related depositional settings. The carbonate bank is composed of fusulinid and algal biomicrites and is flanked by carbonate talus apron deposits dipping at 2–51. These interfinger with delta-fed turbidite systems deposited as slope aprons. Channel and lobe complexes can be distinguished, and form the framework of the prograding aprons. Distally, the sandstones interfinger with dark, basinal shales that extend far into the Midland Basin. The carbonate bank locally stood above the shelf floor, as indicated by the presence of oolitic patches, and dipping flank beds. Narrow clastic-filled channels pass across the shelf edge, and appear to be erosional in origin. The shelf margin is therefore a zone of sediment bypass, and the erosional sandstone-filled channels are interpreted as the heads of shelf-margin canyons that fed clastic detritus to the slope. Erosion of the shelf, deepening of the canyons, and sediment delivery to the slope were enhanced during times of low sea level, while carbonate bank sedimentation was facilitated by episodes of high sea level, when clastic, deltaic sediment input was shifted east and north by marine transgression. Galloway and Brown (1973) also emphasized the importance of autogenic shifting of deltaic distributaries in the generation of the resulting, complex stratigraphy. One of the more economically important clastic units in the Pennsylvanian–Permian succession is the Bartlesville Sandstone, part of the Cherokee Group of Middle Pennsylvanian (Desmoinesian) Age. The Cherokee Group comprises a mixed carbonate–clastic succession reaching 610 m in thickness in Eastern Oklahoma. An early oil discovery in the Bartlesville Sandstone, the Cushing field, in Oklahoma, is estimated to contain 470 million barrels of oil (ultimate recovery). Initially, the reservoirs were described as ‘‘shoestring’’ sands (Weirich, 1929; Bass, 1936). Later, Visher (1968) and Visher et al. (1971) and Berg (1986) interpreted the unit as part of a large delta complex prograding southward from cratonic (and ultimately Appalachian tectonic) sources in and to the east and north of Missouri (Figure 13). Recent work (Ye and Kerr, 2000) has emphasized the widespread unconformity at the base of the Bartlesville, and reinterpreted the sandstone as the deposit of a large fluvial

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Figure 9 Stratigraphic columns for the main basins and platforms of the Permian Basin area (U. S. Geological Survey Digital Data Series DDS-36; 1996; see Scholle, 2006).

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Figure 10 Correlation of electric logs across the margin of the shelf in Scurry County. Texas by Van Siclen (1958), showing the di¡erentiation into the ‘‘three critical environments of deposition’’of Rich (1951). AAPG r 1958, reprinted by permission of the AAPG whose permission is required for further use.

complex within an incised paleovalley. The valley fill ranges from 43 to 85 m thick and is nearly 100 km wide at its southern preserved limit. It consists of braided-fluvial facies in an interpreted lowstand systems tract at its base, and a meandering channel-fill sandstone and floodplain mudstone succession comprising a transgressive systems tract above (Figure 14; Ye and Kerr, 2000). Another important oil-producing clastic unit is the Lower Pennsylvanian Morrow Formation, which occurs on the north flank of the Anadarko Basin along a belt that extends northwest into Kansas and Colorado. This nonmarine to marginal-marine, predominantly sandstone unit, was deposited mainly as the fill of incised valleys during glacioeustatic changes in sea level (Rascoe and Adler, 1983; Bowen and Weimer, 2003).

5. Ouachita Deformation and Sedimentation Largely buried beneath the Gulf Coastal Plain (Galloway, Chapter 15) is the Ouachita Orogen, the southernmost segment of the Appalachian Orogen of Eastern North America, which was formed during the final assembly of Pangea in the Late Paleozoic. Excellent descriptions of the Ouachita Orogen, its stratigraphy, sedimentology, structural geology, and tectonic history, have been provided by Arbenz (1989) and Keller et al. (1989), and in several chapters of Hatcher et al. (1989), which contains a useful summary by Viele and Thomas (1989). Poole et al. (2005) document the continuation of the Ouachita Orogen into Northern Mexico, where it is named the Sonoran Orogen. The largest area of exposure of the orogen comprises the Ouachita Mountains, which straddle the Oklahoma–Arkansas border. A small but important portion of the Ouachita Orogen is also exposed in the Marathon Uplift of Southwest Texas (Figure 1). Extensive petroleum exploration and deep seismic reflection profiling (Lillie et al., 1983) have served to define the subsurface portion of the orogen beneath the coastal plain, although details are sketchy because of deep burial below the Mesozoic cover (Nicholas and Waddell, 1989; Thomas, 1989). Only 100 km southeast of the Llano Uplift, the subcrop of Ouachita rocks beneath the Gulf Coastal Plain is already at a depth of 20 km (Worrall and Snelson, 1989). In the Mid- to Late Mississippian, the progressive collision between Laurentia and Gondwana and intervening terranes that had been underway since the Ordovician (See Chapter 1) reached the southern margin of North

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Figure 11 Van Siclen’s (1958) models of cyclic sedimentation under varying conditions of sea-level change and sediment supply. AAPG r 1958, reprinted by permission of the AAPG whose permission is required for further use.

America (Arbenz, 1989; Thomas, 2006; Ettensohn, Chapter 4). The initial collision is thought to have occurred near the southern tip of the Appalachian Orogen in what is now Central Mississippi (Thomas, 1989, Figure 3C). The colliding arc is interpreted as a series of amalgamated Yucatan and Hunic Terranes, lying between Laurentia and the main Gondwana Continent, within the remnant Rheic Ocean (Figure 11, Chapter 1). Collision occurred along a south-dipping subduction zone, and was essentially complete by the Mid-Pennsylvanian (Figure 12, Chapter 1). The result is a classic collisional foreland orogen, comparable to the Appalachian Belt of the Eastern United States (Figure 15; see also Ettensohn, Chapter 4). Foreland basins of proforeland type (basins resting on the downgoing plate) lie adjacent to the orogen, and include, from west to east, the Val Verde, Kerr, Fort Worth, Arkoma, and Black Warrior basins (Figures 1, 16). These foreland basins overlie the Cambrian–Mississippian rocks of the continental platform margin, which developed a series of extensional faults as the margin was progressively loaded by the developing Ouachitan Orogen. Many of these functioned as growth faults during flysch sedimentation. The Oklahoma ‘‘aulacogen’’ was reactivated as a foreland basin during the same regime of collisional tectonism, with the development of a series of reverse faults extending WNW–ESE across Southern Oklahoma. Uplift on these faults gave rise to the Wichita Uplift and, further west, the Amarillo Uplift. Displacement on these faults may reach 12 km, and resulted in the erosion and transportation of very substantial volumes of clastic detritus northward into the Anadarko Basin, where the Pennsylvanian–Early Permian section, alone, is about 7.5 km thick (Johnson et al., 1988; Arbenz, 1989), including a ‘‘granite wash’’ derived from the Wichita– Arbuckle Uplift to the south. This thins westward to about 2 km in Western Oklahoma (Figure 17). The

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Figure 12 1973).

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The pattern of shelf margin ‘‘reciprocal sedimentation’’ in North Central Texas (modi¢ed from Galloway and Brown,

Pennsylvanian configuration of the basin is essentially that of a foreland basin, the uplift to the south corresponding to a fold-thrust belt created during the transpressive Ouachita Orogeny (Brewer et al., 1983). This is one of the thickest sections of Phanerozoic strata preserved within the North American continent (Johnson et al., 1988). In the core of the Ouachitan Orogen, and now exposed in the Broken Bow Uplift in the Ouachita Mountains (Figure 15), is the assemblage of metamorphosed pre-orogenic, Lower, and Middle Paleozoic rocks that were deposited within the Rheic Ocean and along the Laurentian continental margin (Arbenz, 1989; Nicholas and Waddell, 1989; described briefly earlier in this chapter). This has been termed the ‘‘Interior zone’’ of the orogen (Flawn et al., 1961). The rocks consist mainly of dark green and black shales, thin-bedded siltstones, and turbidite sandstones. They are intensely deformed and cleaved and have undergone low-grade metamorphism. The shales display high organic contents, where not obscured by metamorphism, and probably served as one of the major source beds for the hydrocarbons of the Permian Basin. Although the Ouachitan Orogen is interpreted as the product of continental collision along a subduction zone, there is surprisingly little evidence of igneous activity. Rhyodacite tuffs of Early Mississippian Age have been mapped near the southern margin of the exposed Ouachita Mountains (Arbenz, 1989), and a rhyolite porphyry with an overlying tuff, of probably Carboniferous age have been encountered in deep wells drilled into the basement near the Texas–Louisiana border (see Figure 1) (Nicholas and Waddell, 1989). Sedimentation rates indicate that, commencing in the Mid- to Late Mississippian, the Ouachita Orogen and its foreland began quite suddenly to undergo rapid subsidence (Figures 18, 19). This can be interpreted as a result of the downbending of the continental margin caused by continental collision and resultant crustal thickening, the process well described by Stockmal et al. (1986). The change in sedimentation rate was not simultaneous everywhere, but took place earliest in the Oklahoma ‘‘aulacogen’’ (or foreland basin); later in the Arkoma Basin. The most rapid sedimentation rates were during the Mississippian and Early Pennsylvanian in the Ouachita Mountains area, and during the Mid-Pennsylvanian (Atokan) in the Arkoma Basin (Figure 18). Sedimentation rates calculated by Sloss (1988) for the Absaroka I sequence (Latest Mississippian to Early Permian) exceeded 100 m/Myr in the Marathon area, and in the Fort Worth and Anadarko foreland basins, and in the Arkoma Basin they reached 1 km/Myr. Sedimentation styles changed dramatically at the same time as sedimentation rates increased, with the appearance in the foreland of a sedimentary assemblage long referred to as ‘‘flysch.’’ This term, applied initially to a syntectonic, deep-water sedimentary unit in the Swiss Alpine foreland basin, has been widely used for comparable sedimentary–tectonic assemblages in other basins since the 1930s. In the Ouachita Basin 6 km of deep-water flysch sediments were deposited during the Mid-Mississippian to Early Pennsylvanian (Stanley and Jackfork Groups), and there are 9 km of Atokan (Mid-Pennsylvanian) flysch in the Arkoma Basin (Figure 20;

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Figure 13 The extent of the Bartlesville Sandstone ‘‘delta’’ in Kansas and Oklahoma (Berg, 1986). This system £owed axially along the Alleghanian foreland basin of the Appalachian system, and was £anked to the west by a carbonate platform situated over the associated forebulge (see Figure 27, Chapter 4).

Arbenz, 1989). A maximum of 4.2 km of Pennsylvanian flysch is present in the Marathon area. Flysch sedimentation continued later in this most southwesterly area of the orogen, ending in the Earliest Permian (Ross, 1986; Meckel et al., 1992). The predominant sedimentary facies in these thick flysch successions is sandstone–siltstone–shale turbidities and related sediment-gravity flow deposits, deposited in various submarine fan environments (Figure 21; Morris, 1989). There are also numerous beds of pebble- to boulder conglomerate and olistostromes, with blocks, boulders, and slabs up to several hundred meters across. These originated as slumps on the continental margin, the slope of which was rendered unstable by the tectonic loading. Some of the olistostrome units show structural disturbance, which has been interpreted as the product of syndepositional tectonism in the active, advancing front of the orogen (Morris, 1989). A similar transition from passive-margin continental slope sedimentation to flysch took place somewhat later in the Marathon area (Ross, 1986; Wuellner et al., 1986; Meckel et al., 1992). Here, tectonism, in the form of

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Figure 14 A model of Bartlesville Sandstone in the form of generalized gamma ray logs and a stylized longitudinal cross-section through Eastern Oklahoma. The section reaches about 70 m in thickness towards the south (Ye and Kerr, 2000). AAPG r 2000, reprinted by permission of the AAPG whose permission is required for further use.

Figure 15 Structural cross-section through the Ouachita Mountains. Adapted from Arbenz (1989).

folding and thrust faulting, is dated as Late Pennyslvanian to Early Permian (Late Desmoinesian to Wolfcampian). A flysch-type turbidite assemblage with the same age range was deposited in the adjacent Val Verde foreland basin. Facies trends and a wealth of paleocurrent data from the outcrop areas indicate generally westward transport directions in the Ouachita Mountains area, parallel to the continental slope (Figure 21; Houseknecht, 1986), with some smaller marginal fans deposited on the continental slope from sediment-gravity flows flowing ‘‘transversely’’ (southward) down the continental slope. In the Marathon area some units are similarly ‘‘axial’’ in transport direction (northeast to southwest), while other units show southwesterly (‘‘transversely’’ off-shelf) and northwesterly transport directions. These tectonic, facies, and paleocurrent data have been interpreted in terms of a ‘‘remnant ocean basin’’ model by Graham et al. (1975) and Ingersoll et al. (1995). In this model, uplift, erosion, and the generation of detritus occur primarily along an active collisional orogen, with transport taking place into the nearest low-standing depocentres. Along an orogen closing by oblique collision, the remaining area of open ocean becomes the most important depocentre, with sediment transported into it parallel to tectonic strike. Sedimentation is diachronous, extending in front of the advancing orogen as the ocean continues to close along

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Figure 16 Isopachs of Pennsylvanian strata (adapted from Cook and Bally, 1975). Major Ouachitan foreland basins are also shown.

strike. The Bengal Basin between India and Burma (Myanmar), has long been regarded as the type example of a remnant ocean, and the Ouachitan Basin as the type example of an ancient analog (Graham et al., 1975). The Ouachitan Orogeny commenced with growth faulting on the continental slope in front of the Ouachitas in the Mid-Pennsylvanian (Atokan), indicating the beginning of thrust sheet loading (Arbenz, 1989). Movement on thrust faults and folding in the central and southern Ouachitas commenced at about the same time (Figure 22). As obduction of the Ouachitan allochthon continued, deformation extended further northward towards the foreland, reaching its eventual limit in the Late Pennsylvanian (Mid-Desmoinesian). A classic thin-skinned fold-thrust belt was generated, with a major de´collement above the downflexed platform sediments, and stacked, north-verging thrust faults carrying slices of Upper Mississippian to Mid-Pennsylvanian flysch northward towards the foreland (Figure 15). Tectonism and flysch sedimentation began later (Late Pennyslvanian–Desmoinesian) and ended later in the Marathon area (Meckel et al., 1992), consistent with the model of progressive westerly closure of the remnant ocean basin. A clear indication of the termination of orogenic deformation is indicated by an angular unconformity between rocks of Wolfcampian and Leonardian (Early Permian) Age in the Marathon inlier (Ross, 1986). Little is known about the Ouachita Orogen beneath the deeper parts of the Gulf Coastal Plain. Wells drilled into basement in Arkansas and Mississippi have encountered slates with well-defined cleavage and containing abundant quartz veins. Limited seismic and well data show that the broad antiform of the Broken Bow Uplift (Figure 15) continues southwestward beneath Texas as the Waco Uplift, which is located immediately to the southeast of a 32–65-km-wide frontal thrust zone (Nicholas and Waddell, 1989). The Devils River Uplift in the Marathon area (Figure 2) is cored by Precambrian rocks and may be an exposed portion of a similar interior uplift (Nicholas and Waddell, 1989). The same broad structure curves southeastward from the Ouachita Mountains into the subsurface of Southern Arkansas. Beneath East Central Mississippi, deformed and metamorphosed slates are cut by thrusts showing Appalachian (NE–SW) trends (Thomas, 1989). All these uplifts are interpreted as ‘‘external massifs of the Ouachita Orogen, which formed late in the collisional cycle’’ (Nicholas and Waddell, 1989, p. 668). High levels of thermal maturation of the organic residues and a regional level of greenschist metamorphism are tentatively attributed to this final stage of collision (compare to the Southern Appalachians; Ettensohn, Chapter 4).

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Figure 17

Stratigraphic cross-section through the Anadarko Basin Bally (1989). Adapted from Adler et al. (1971).

Several deep wells drilled into the pre-Coastal Plain basement of the Texarkana Platform and further south, near the Texas–Louisiana border have encountered relatively undeformed sediments of Middle Pennsylvanian (Desmoinesian) Age. They are mainly carbonate sediments of shallow-marine shelf origin. In one location they overlie a succession of rhyolitic volcanics (location shown in Figure 1) which, in turn, are thought to unconformably overlie deformed Mid-Pennsylvanian Age (Atokan) flysch (Nicholas and Waddell, 1989). The presence of these rocks has been interpreted in terms of a late-orogenic, ‘‘successor’’ basin overlying the Ouachitan Orogen. The Desmoinesian Age of these rocks postdates the uplift of the orogen in the Ouachita Mountains area, but is contemporaneous with the sedimentation and fold-thrust tectonism of the Marathon area and with the syntectonic clastic–wedge sedimentation of the Anadarko Basin. The development of a carbonate platform on the eroded roots of a still active orogen is comparable to some areas of modern Indonesia and the Philippines.

6. The Permian Basin and the Capitan Reef The platform and basin topography initiated by Ouachitan deformation during the Mississippian became fully established during the Pennsylvanian as a consequence of the continuing tectonism (King, 1948), although the classic Permian Basin basinal configuration and the contained rocks described in this section, are entirely postOuachitan in terms of their tectonic development. Reactivation of lineaments in the Older Paleozoic basement (which we now interpret as having developed during the breakup of Rodinia), may have guided some of the basin and platform differentiation during the Later Paleozoic tectonism (Figure 23; Hills, 1984). On the Matador Uplift and on parts of the Central Basin Platform and the Diablo Platform, Pennsylvanian or Permian strata now rest on Precambrian Basement (Frenzel et al., 1988). These structures were further accentuated by differential sedimentation and regional subsidence during Permian time, as can be reconstructed from the preserved

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Figure 18 Subsidence curves for the Ouachita orogen and foreland (depths in km). Inset shows the relationship between the slope of the subsidence curve and sedimentation rate, in m/Myr (Arbenz, 1989).

architecture of platform and basin and the clinoform stratigraphy that links the two. In this scenario, as suggested by Scholle (2006), original structural relief was significantly accentuated by higher rates of sedimentation of shallow-water carbonate deposits on structural ‘‘highs’’ compared with lower rates on structural ‘‘lows.’’ Reef bodies partially or entirely encircled the basinal areas (e.g., Figure 24). Basins which may have been only a few tens of meters deep at the start of the Permian eventually had water depths in excess of 500 m by the close of MidPermian (Guadalupian) time. Sedimentation rates in the Delaware and Midland Basin were exceptionally high (W120 m/Myr) during the deposition of Absaroka II subsequence (Middle–Upper Permian) (Sloss, 1988). Isopachs for the total Permian section show a belt of maximum thickness extending northward across West Texas, with the thickest section occurring in the Delaware Basin (Figure 25). Formal stratigraphic columns for the main basins and platform areas in the Permian Basin are shown in Figure 9. A paleogeographic reconstruction of the Permian Basin during Leonardian (Early Permian) time is shown in Figure 24. King’s (1948) classic reconstruction of the Guadalupian (Mid-Permian) Capitan Reef is reproduced as Figure 26. A modern stratigraphic synthesis of the Delaware Basin is provided as Figure 27 and the spectrum of facies occurring in the Guadalupe sections is shown in Figure 28. Pennsylvanian strata are characterized by rapid lateral facies changes and are highly variable. On the Northwest Shelf the deposits are primarily clastic (Morrow and Atoka Series, up to 460 m total thickness) comprising a partially nonmarine succession of fluvial channel and incised valley-fill deposits, similar to the older Cherokee Group of Oklahoma and deposited under similar, cyclic conditions. To the south, on the Diablo Platform, Pennsylvanian strata are primarily fossiliferous limestone, deposited in a marine carbonate shelf environment. These pass into fine-grained terrigenous clastics in the Delaware and Midland basins, a thin succession in the latter indicating somewhat starved conditions. The basins were not yet actively subsiding, and the Midland Basin is notable for the presence of a large reef complex, the Horseshoe Reef, which began accumulating in the Mid-Pennsylvanian, and continued into the Early Permian (Wolfcampian) (Figure 24). The following account of Permian stratigraphy and sedimentology of the Permian Basin is based largely on Frenzel et al. (1988) and Scholle (2006). Paleogeography during the Early Permian (Wolfcampian) is shown in Figure 24. By this time the basin and shelf topography of the Permian Basin had become fully established, with a paleogeographic differentiation into reef, back-reef, fore-reef, and basin similar to that of the modern Bahamas Platform. An 800 m thick section of mainly reefal Pennsylvanian–Permian carbonates define the eastern margin of the Permian Basin and the structural boundary with the Eastern Shelf. These show the high-frequency cyclicity discussed in the previous section.

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Figure 19 Contrasting patterns of subsidence in the Ouachita Basin and in foreland basins. Pie diagrams for the Arkoma Basin indicate the mount of time represented by each of the three genetic packages of strata, expressed as a total of the Paleozoic time recorded by the strata (after Houseknecht, 1987).

The most well known part of the stratigraphy (because of the classic outcrops in the Guadalupe and Delaware Mountains) is the Leonardian–Guadalupian succession. A coastal fluvial and eolian succession of detrital redbeds (Abo Formation) occurs in the northern part of the shelf. These sediments were probably derived by erosion of remnant basement uplifts of the Ancestral Rockies. They consist mainly of sandstone and siltstones with evaporite nodules and beds of gypsum. To the south, these beds pass into dolmicrites containing pellets, oncoids, calcispheres, and some foraminifera and algal stromatolites, deposited in a coastal, hypersaline lagoon. These are variously referred to the San Andres Formation or the Artesia Group (Figure 27). A distinctive, unusual bed of pisolitic dolomites, with polygonal ‘‘teepee’’ structures lies to the south of the lagoon facies, and has been the subject of some controversy regarding its origins. The currently most plausible hypothesis is that these beds represent a barrier deposit that has been subject to substantial penecontemporaneous alteration as a result of paleosoil (caliche) development and marine seepage, resulting in mineralogical replacement and recrystallization. The shelf-margin facies consists of a wide variety of marine limestone lithologies. Most of the beds are grainstones or packstones, together with some wackestones and algal boundstones. The rocks are richly fossiliferous, containing fusulilinids and other foraminifera, gastropods, pelecypods, and algae. These beds, together with the pisolitic beds behind them, formed a seaward-facing barrier, and were probably formed within a mosaic of biostromal islands and bioclastic lenses, with abundant small tidal passes. The shelf margin rocks pass seaward into the main reef facies, constituting the Goat Seep Dolomite and the overlying Capitan Limestone (Figure 27). The Goat Seep Dolomite consists of massive to thick bedded dolomitic limestone. The Capitan Reef zone rocks are composed of an organic framework structure displaying considerable faunal diversity. The main framework-building organisms are calcareous sponges and phylloid algae, and these occur with echinoids, brachiopods, bryozoans, mollusks, ostracods, scarce solitary corals, and trilobites. Sponge-algal rubble interfingers with algal grainstones on the shelf side of the reef. There is substantial evidence for contemporaneous alteration and cementation of the deposits. Facies relationships and the nature of the early diagenetic modifications of the rocks indicate that the Capitan undoubtedly was an actual topographic reef during its formation.

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Figure 20

Generalized stratigraphic columns in the Ouachita Basin and in foreland basins (Arbenz, 1989).

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Figure 21

Conceptual, schematic paleogeographic diagram for the Ouachita Basin (Houseknecht, 1986).

Figure 22

Timing of initial tectonic development of the major tectonic elements in the Ouachitan Mountains (Arbenz, 1989).

A

WARD CO.

A'

CRANE CO.

+ 1000 YATES, SEVEN RIVERS AND QUEEN FMS.

- 1000

SAN ANDRES FM.

- 2000

CLEARFORK FM.

- 3000

WICHITA - ALBANY FMS.

- 4000

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(Ft) + 2000

- 5000 - 6000

- 8000

F MP

M.

A

LFC

ROUS

ONIFE

CARB

NIAN DEVO RIAN U IL S

WO

- 9000 - 10,000 - 11,000

1000ft IA DOVIC

R

N-O

BRIA

CAM

2000ft

N

WARD CO.

A

A′ CRANE CO.

0

3000ft 6000ft

- 7000

- 12,000 - 13,000

Figure 23 Structural cross-section through the Guadalupe Mountains (Scholle, 2006; based on King, 1948, Hills, 1984 and other sources). Note that the Capitan Reef margin developed above a major fault in the Older Paleozoic basement.

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Figure 24 Paleogeographic reconstruction of the Permian Basin during the Leonardian (Early Permian). The reef crests are composed of the Wichita--Albany Formation (after Wright, 1979).

In front of the reef is the slope facies of the Capitan Limestone. This is a steeply dipping fore-reef talus apron (dips W351 at the top of the slope, decreasing basinward) composed of rubble derived from the reef and back-reef rocks. Brachiopods, echinoderms, and bryozoans appear to have lived in this environment, which was periodically swept by submarine landslides, and carbonate debris flows and turbidity currents. Occasional siliciclastic– sandstone filled channels up to 10 m deep occur within the slope facies. These probably represent the main paths by which terrigenous detritus bypassed the shelf margin and reef to be deposited ultimately in the basin. At the foot of the slope, bed thicknesses decrease, siliceous biota and siliceous replacement of fossils become more common, and the detrital carbonate units thin out to comprise no more than 10% of the succession in the basin. Most of the basin fill consists of sub-feldspathic very fine-grained sandstones and siltstones, similar in composition to the clastic deposits of the shelf. These have been assigned to the various units of the Delaware Mountain Group (Figure 27). Facies relationships and sedimentary structures indicate that these deposits were formed as submarine channel and levee deposits of detrital aprons (Harms, 1974; Williamson, 1977). The shelf to basin spectrum of facies described earlier interfinger and alternate in a cyclic pattern (Figure 28). At least ten major tongues of slope facies interfinger with ten tongues of Delaware Mountain Group basinal clastics, through the Guadalupian Stage, which represents about 10 Myr of the Mid-Permian. The Capitan Reef complex therefore exhibits the same pattern of cyclic sea-level change as that displayed by other Permian Basin deposits, as summarized in the previous section. The model of reciprocal sedimentation, described earlier, has also been invoked to explain the patterns of sedimentation there, with reef accumulation and talus-slope progradation occurring more rapidly during times of high sea level, and tidal-channel and canyon incision being enhanced during times of low sea level, when the bulk of the terrigenous detritus was probably transported across the shelf and out into the deep basin. Regional tectonic events, including epeirogenic tilting and variations in intraplate stress reflecting variations in sea-floor spreading rates undoubtedly also affected relative sea levels.

7. Evaporite Sedimentation in the Delaware and Midland Basins The Ouachitan Orogen formed a barrier to the south of the Permian Basin at least as early as the Early Pennsylvanian. A connection with marine environments to the west was maintained through a narrow seaway called the Hovey Channel, located to the north of the Marathon Uplift (Figures 1, 2). This marine connection

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Figure 25 Permian isopachs (after Cook and Bally, 1975).

Figure 26 King’s (1948) classic stratigraphic reconstructions of the Capitan Reef of the Guadalupe Mountains, across the northern margin of the Delaware Basin.

became increasingly restricted during the Early Permian, probably partly as a result of falling global sea levels. Saline brines began to form, first in the marginal parts of the Permian Basin and then, progressively, throughout the entire basin. Evaporite sediments, initially anhydrite and then halite, began to accumulate in the Palo Duro Basin during the Leonardian (Wichita and Clear Fork Groups of Leonardian Age and Lower San Andres Formation, which is Guadalupian). Salt precipitation began in the Midland Basin during the Guadalupian. The thickest salts are generally observed on the parts of the shelf away from the Delaware Basin toward the east and north. Several major cycles of sandstones, anhydrite, and halite of the Yates Formation were deposited across the

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NORTHWESTERN SHELF

Dol.

Cherry Canyon Ss. Lo. San Andres / Cutoff Fm.

Cutoff Fm.

Bell Canyon Ss. Cherry Canyon Ss. Brushy Canyon Ss.

Ochoan Guadalupian

G a si Ar te

Guadalupian

meters 1500

Foreslope

Seep

Victorio Peak Dol.

Castile Fm.

Capitan Ls.

Goat

Up. San Andres

Leonardian

1000

e

ssiv

Ma

Queen

Grayburg

SE

Salado Fm.

Tansill Yates Seven Rivers

p.

0

500

DELAWARE BASIN

Salado Fm.

Delaware Mtn. Gp.

NW

Cutoff Fm. Bone Spring Ls. 1

0

2 mi Vert. Exagg. ca. 3x

0

1

2

Approximate time line

3 km

Figure 27 Modern stratigraphic synthesis and nomenclature of the transition from the Northwestern Shelf into the Delaware Basin (Scholle, 2006).

Redbeds & sabkha/ salina evaporites

Pisolitic & grainstones

Reef

Lagoonal mudstones

Back-reef grainstones

Forereef

Basin

Ca. 1 km

Figure 28 Facies relationships in the Capitan Reef complex (Scholle, 2006).

platform during a sea-level lowstand; the corresponding deposits in the Delaware Basin are in the Bell Canyon Formation. The deposits of the following highstand, also composed of a number of cycles, are carbonate, anhydrite, halite, and sandstone of the Tansill Formation. The Lamar Limestone at the top of the Bell Canyon Formation is the basinal equivalent to the Tansill. By the end of the Guadalupian, sedimentation had mostly leveled the shelf topography east of the Delaware Basin, so that the major structural elements such as the Central Basin Platform, Midland Basin, Northern Shelf, Matador Arch, Eastern Shelf, and Ozona Platform were expressed only by subtle contrasts in subsidence rates. Stratigraphic units extend across structural positive areas with only minor changes in thickness or composition. Post-Guadalupe sedimentation is summarized by Frenzel et al. (1988) and Scholle (2006). The remnant Delaware Basin was rapidly filled by massive evaporites during Late Permian (Ochoan) time. The first phase of filling was the Castile Formation, consisting of millimeter-scale interbedded laminae of gypsum (anhydrite in the subsurface) and organic matter plus calcite. The laminae are planar and highly continuous, having been correlated over much of the Delaware Basin (Anderson et al., 1972). Approximately 260,000 lamination cycles have been counted and these have been interpreted as varves, with the gypsum representing evaporative summer–fall layers and the organic matter plus calcite reflecting more normal-salinity winter–spring conditions. It has also been

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suggested that the laminae reflect Milankovitch-scale climatic variability including precession (ca. 20,000 year) and eccentricity (100,000 year) cycles (Anderson, 1984), although we cannot be certain of the nature or periodicity of orbital parameters so far in the geologic past. This combination of characteristics suggests that the Castile evaporites were deposited in water depths of tens to hundreds of meters (at least below wave base). The Castile sediments filled the Delaware Basin and evaporite sedimentation then spread across the surrounding shelfal regions. Water depths were undoubtedly much shallower at this stage and conditions became even more hypersaline with deposition of the anhydrite, halite, sylvite, and other bittern minerals of the Salado Formation. Commercial deposits of potash minerals have been exploited in these thick, last-stage fillings of the Delaware Basin and its marginal areas. Water depths during this phase were probably quite shallow, perhaps of the order of a few meters or less (Lowenstein, 1988). Subsequent to formation of the Salado evaporites, sedimentation in the area was dominated by nonmarine redbed deposition which extended into the Triassic Period, completing the cycle of progressive desiccation and filling of the Permian Basin region. These evaporites preserved the geometry of the basin and provided the excellent seals needed to trap a high percentage of the oil generated from the rather sparse basinal source rocks.

8. Oil and Gas Production The first commercially successful oil well drilled in what was to be called the Mid-continent Oil Field was the Norman No. 1 near Paola, Kansas, on November 28, 1892. The successes that followed the Nellie Johnstone No. 1 (Bartlesville, Oklahoma) discovery in 1897, demonstrated the existence of a large oil field in the Central and Southwestern United States. It became known as the Mid-continent Oil Field. Continued drilling found many other oil fields and pools within the Mid-continent, both large and small. The Glenn Pool, discovered in 1905, was not the biggest oil discovery in Oklahoma, but for a time, it was the nation’s largest producing oil field. The discovery well flowed at about 85 barrels a day. The Cushing Field was discovered in 1912 and the Oklahoma City Field was discovered in 1927. For a time, the Oklahoma City Field was the nation’s largest oil field. All of these fields produce from the Pennsylvanian Bartlesville Sandstone. Oklahoma remains the US state with the fifth largest remaining reserves, at 588 million barrels (93 106 m3). The Deep Anadarko Basin of Western Oklahoma is one of the most prolific gas provinces of North America. Wells drilled here have been among the world’s deepest. The Bertha Rogers No. 1 in Washita County, drilled in 1971 to 9583 m (31,441 ft), was then the world’s deepest well. In 1979 the No. 1 Sanders well near Sayre became Oklahoma’s deepest gas producer at 7619 m (24,996 ft). Panhandle-Hugoton, the largest North American gas field produces mainly from Permian reservoirs, but is probably sourced from much deeper. Drilling to the Lower Paleozoic of the Permian Basin in Texas began in the 1920s, and extensive exploration began in the Permian Basin in the 1930s. Petroleum has been produced in the Permian Basin from most units, from the Cambrian to the Cretaceous, although the bulk of the production is from the Permian section, with post-Permian production negligible (Dolton et al., 1979; Scholle, 2006). In the Permian Basin 91.6 billion barrels (14.5 109 m3) of oil and 106 trillion cubic feet (3 trillion cubic meters) of gas have been discovered. Some 65 billion barrels (10.3 109 m3) of oil in place, representing 71% of the known total, is located in Permian reservoirs, in more than 2,000 separate pools. Two-thirds of this is from reservoirs of Guadalupian Age, most of the remainder in the Permian is in Leonardian-age reservoirs (Dolton et al., 1979; Scholle, 2006). Traps in the Permian are typically combination structural-stratigraphic traps, the multiple lateral and vertical facies changes and the many widespread porosity–permeability changes generated by diagenesis providing multiple trap configurations. Evaporites, which increase in importance up-section, provide the sealing units in many cases. Dolomites of the San Andres Formation and back-reef sandstone and dolomite of the Yates, Seven River, and Queen formations constitute the major reservoir units. Another, smaller set of pools is located in channel sandstone and basinal limestones of the Delaware Mountain Group and Bone Springs Formation. Reef rocks of the Capitan, Goat Seep, and other units are not productive, owing to early diagenetic porosity reductions. The major source rocks are organic carbon-rich basinal sediments of the Bone Springs Formation and the Delaware Mountain Group. Oil and gas production from the Ouachita belt, including its foreland basins, is very modest, by comparison with the Permian Basin proper. Most traps are structural, such as those related to roll-over above thrust faults. Meckel et al. (1992) reported the discovery of a total of 500 million barrels (7.95 107 m3) of oil and 10 trillion cubic feet (2.83 1011 m3) of gas.

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Figure 29 Schematic paleogeographic sketches.

9. Summary The evolution of the southern part of the continent can be summarized in terms of four broad stages (Figure 29): (A) From Late Precambrian to Mid-Mississippian the area constituted a passive continental margin with a stable craton to the north. The breakup of the supercontinent Rodinia generated rifts that penetrated the craton at high angles, and became the basis for two major, persistent sedimentary basins, the Reelfoot Rift underlying the Mississippi Embayment, and the Okalahoma ‘‘aulacogen’’, the precursor of the Anadarko Basin. Shallow-water carbonate and clastic sediments accumulated on the craton, and deeper water clastics, including sediment-gravity flow deposits, on the continental slope; deposits now preserved as deformed units within the Ouachita Orogen. (B) From the Late Mississippian to the Early Pennsylvanian the continental margin underwent progressive, diachronous collision with an arc complex comprising various amalgamated terranes. A carbonate bank persisted in places in the craton, but shallow-water, mixed sand–mud sedimentation was widespread. Structural disturbance of the continental margin led to occasional margin collapse and olistostrome development, while turbidites and related sediment-gravity-flow deposits began to fill the remnant ocean basin lying to the south. Deltaic complexes were shed from the rising collisional orogen in Mississippi and Alabama.

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(C) The Ouachita Orogeny climaxed during the Mid- to Late Pennsylvanian. The continental margin was depressed by the collisional load and underwent down-to-the-south extensional and growth faulting. A series of proforeland basins developed in front of the growing fold-thrust belt, and filled with many kilometers of flysch-type clastics, dominated by repeated turbidity-current deposits. The same transpressive collisional orogeny generated the Ancestral Rocky Mountains to the west, and these served as a barrier to circulation and as a sediment source during the remainder of the Paleozoic. In-plane stresses from the orogeny acting on the craton accentuated existing structural weaknesses and began to differentiate the basinand-shelf architecture which formed the basis for the subsequent development of the Permian Basin. The Oklahoma Aulacogen was transformed into a basin of foreland type, bordered by an active fold-thrust belt to the south. The combination of early rifting and subsequent foreland-basin development gave the resulting Anadarko Basin, at 12 km, the thickest sediment fill of any Phanerozoic basin in North America. Sedimentation across the craton throughout the Pennsylvanian was characterized by the formation of cyclothems. These have long been interpreted as the product of high-frequency cyclicity, generated by glacioeustasy. (D) The classic Permian Basin is characterized by a well-developed cratonic shelf and basin architecture which was accentuated by the style of Early to Mid-Permian sedimentation. On the structurally more elevated platform areas reef carbonates developed, while slower sedimentation of deep-water clastics was underway in the basinal areas. Much of the continental interior became uplifted and exposed during the Permian, but marine circulation was maintained by a connection to the ocean through the Hovey Channel, to the southwest. However, this became progressively restricted during the Permian, and in the Late Permian sedimentation was dominated by the development of thick evaporite deposits.

ACKNOWLEDGMENTS The author is very grateful to Peter Scholle and the New Mexico Bureau of Geology and Mineral Resources for permission to use several of the diagrams posted on his website. The chapter was reviewed by Bill Galloway and Ron Blakey, who provided many useful comments.

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CHAPTER 9

The Western Interior Basin Andrew D. Miall, Octavian Catuneanu, Boyan K. Vakarelov and Ryan Post

Contents 1. Introduction 2. Geodynamic Framework 3. Paleogeographic Evolution 3.1. Jurassic 3.2. Early Cretaceous 3.3. Late Cretaceous 4. Allogenic Mechanisms of Sequence Development 4.1. Tectonic cyclicity 4.2. Basement control 4.3. Milankovitch cyclicity 4.4. Discussion 5. Economic Resources 5.1. Oil and gas 5.2. Coal 6. Conclusions Acknowledgments References

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Abstract The Western Interior Basin extends north–south over about 351 of latitude, from Texas to the Northwest Territories, a distance of more than 3,000 km. The basin developed as a result of crustal loading during the western migration of the North American Plate and the subduction of Panthalassa. Initiation of the Western Interior Basin as a distinctive geodynamic and stratigraphic province is traditionally associated with the deposition of the Upper Jurassic (Oxfordian– Tithonian) Morrison Formation in the United States, and the Fernie and Kootenay Formations in Canada. Crustal loading occurred as a series of pulses, as successive terranes arrived at and were obducted onto the western Laurentian margin. These events are represented in the basin as a series of clastic wedges. Westerly sediment sources associated with contractional tectonism appeared for the ﬁrst time in the Late Jurassic, including the Mesocordilleran Geanticline of Nevada. The ﬁrst major clastic wedge, constituting the Morrison and Kootenay formations, continued into the Berriasian, but much of the Berriasian to Barremian (Neocomian) interval is represented by a regional unconformity throughout the Western Interior Basin. This period corresponds to a ‘‘magmatic lull’’ in the Cordillera. The base of the Cretaceous section, of Late Berriasian or Aptian age, typically consists of a sheet of coarse, ﬂuvial gravels, throughout much of the Western Interior Basin. Provenance studies of the foreland-basin strata indicated that following the regional Mid-Cretaceous episode of tectonic quiescence, erosion tapped into oceanic-arc and related rocks, and syndepositional continental magmatic rocks of Quesnellia, far to the west of the orogenic front. Examination of the ages of conglomerates deposited at this time, and reconstruction of the subsidence histories suggest that a new phase of ﬂexural loading and subsidence commenced shortly after deposition, initiating a new ‘‘constructive’’ phase of development of the Cordilleran orogen. At least two Cretaceous cycles of transgression occurred in northern Canada, but marine waters did not extend southward into the Western Interior Basin until the Aptian. During the Aptian–earliest Albian interval, most of the Western Interior Basin was occupied by ﬂuvial and estuarine systems assigned to such units as the Mannville Group in Alberta– Saskatchewan, and the Kootenai Formation of Montana. The Upper Cretaceous stratigraphy of the Western Interior Basin is characterized by the deposits of several major marine transgressions. Gaps in the stratigraphic record are numerous; some represent millions of years, although most are less than one million years in duration. Eustatic sea-level changes were probably partly responsible for this stratigraphic architecture, but regional and local tectonic processes were also Sedimentary Basins of the World, Volume 5 ISSN 1874-5997, DOI 10.1016/S1874-5997(08)00009-9

r 2008 Elsevier B.V. All rights reserved.

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important. During the Turonian the sea reached an all-time high, calculated to be at least 300 m higher than at present. Shallow seas may have extended over much of the Canadian Shield, with a connection through Hudson Bay to the north Atlantic Ocean. High-frequency sequence cyclicity in some parts of the basin suggest a control by orbitally forced eustasy, which caused sea-level changes in the range of 10 m over time scales in the range of 10–125 ka. Termination of the Western Interior Seaway and foreland basin began during the Late Campanian or Maastrichtian throughout much of the United States, as shallow subduction of the Farallon Plate on the western margin of the continent led to the Laramide Orogeny and to the break-up of the basin.

1. Introduction The Western Interior Basin, which was occupied for about 100 million years by the Western Interior Seaway, is one of the largest, best preserved, and most intensively studied of any sedimentary basin in the world. It represents one of the largest foreland basin systems on Earth, extending north–south over about 351 of latitude, from Texas to the Northwest Territories (Figure 1), a distance of more than 3,000 km. Geologically, the basin is deﬁned by its tectonic setting and origin, as a retroarc foreland (retroforeland) basin (Figure 2) lying above a ramp formed by the crust of the Canadian Shield (Figure 3) and the overlying Paleozoic–Early Mesozoic extensional continental margin. The basin was created by loading and depression of the Shield basement by the thrust masses and accreted terranes of the Cordilleran Orogen between the Middle Jurassic and the Early Cenozoic. At times, the width and depth of the basin are greater than can be accounted for by this plate-margin process, and were increased by a ‘‘dynamic loading effect’’ related to mantle convection processes, as discussed below. Crustal loading periodically resulted in subtle reactivation of tectonic elements in the basement (Figure 3), which are now recognized as a suite of intracratonic structures (Figure 4) that inﬂuenced sedimentation at various times throughout the basin history, especially during Laramide time in the United States (see Lawton, this volume). The distinctive paleogeography of the seaway throughout its history represents variations on a basic theme: clastic sediment sources were generated at different times by contractional orogeny at discrete locations along the length of the Cordilleran Orogen. These shed detritus eastward, forming clastic wedges deposited in a range of nonmarine and shallow-marine environments (Figure 2). The basin center was at most times occupied by a marine seaway in which ﬁne-grained clastic sediments accumulated (Figure 2). At times of tectonic quiescence and/or high eustatic sea levels, the marginal clastic wedges were limited in scale, and the basin was dominated by the relatively deep marine deposits of the basin center. At these times clastic sources were so diminished that conditions in the basin center, especially in the south, became suitable for widespread biogenic carbonate sedimentation. At times of tectonic activity and/or low sea level, clastic wedges extended to ﬁll much of the basin, and ﬂuvial drainages coalesced to form large trunk rivers draining along the axis of the basin. This pattern was most clearly expressed during the Aptian–Albian period, when a major northwestward-ﬂowing system developed. The eastern margin of the basin, extending from the Northwest Territories south through northeastern Alberta, Manitoba, the Dakotas, Nebraska, Kansas and Oklahoma, merged imperceptibly with the shallow seas of the craton, the deposits of which now underlie the Great Plains. Sediment sources on this side of the basin were limited, and the deposits consist mainly of ﬁne-grained clastics. Initiation of the Western Interior Basin as a distinctive geodynamic and stratigraphic province can conveniently be associated with the deposition of the Upper Jurassic (Oxfordian–Tithonian) Morrison Formation in the United States, and the Fernie and Kootenay Formations in Canada. The Morrison and Kootenay are largely nonmarine in origin, and contain the ﬁrst evidence of western provenance, implying the ﬁrst emergence of tectonic lands associated with the growth of the Cordilleran magmatic arc and associated tectonic complexes. The Stump Formation (Oxfordian) of Wyoming, Idaho and Utah also was derived from westerly sources (Cross, 1986). Prior to the Oxfordian, basin conﬁguration in the Rocky Mountain states largely reﬂected the tectonic grain associated with the Ancestral Rocky Mountains (Blakey, Chapter 7). In Canada, pre-Fernie strata preserved in the Rocky Mountain foothills are Shield derived (Ricketts, 1989). The Fernie Formation may in part be composed of sediment derived from the incipient foreland basin forebulge. Ross et al. (2005) used U-Pb geochronology of detrital zircon and monazite and Sm-Nd isotope geochemistry to investigate sediment sources for the foreland-basin strata in Alberta, and concluded that the Fernie–Kootenay succession was derived largely from uplift and erosion of miogeoclinal rocks, such as those now exposed in the Rocky Mountains fold-thrust belt. Younger foreland basin strata contain evidence of sources deeper into the Cordilleran orogen, as noted below. Termination of the Western Interior Seaway and foreland basin began during the Late Campanian or Maastrichtian throughout much of the United States, as shallow subduction of the Farallon Plate on the western margin of the continent (Dickinson and Snyder, 1978; Cross, 1986) led to the Laramide Orogeny. This distinctive

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Figure 1 The extent of the Western Interior Seaway, showing the present erosional edge, and reconstructed isopachs for the Upper Cretaceous succession, chosen to illustrate the shape of the basin during a period when it was at its maximum extent (from Beaumont et al., 1993).

orogenic episode was characterized by faulting and the break-up of the foreland basin into a system of smaller, fault-bounded basins (Lawton, Chapter 12), which began to disrupt patterns of subsidence and sediment dispersal as early as the Late Campanian (DeCelles, 2004), with the climax of the orogeny spanning the Maastrichtian to Eocene. Laramide-style tectonism did not occur in Canada or Mexico. Foreland fold-thrust tectonism continued until the Early Paleogene (although the tectonism is commonly referred to as Laramide because of its age, a practice to be avoided), and the pattern of clastic-wedge progradation referred to above was not brought to a close until regional uplift in the Late Paleocene or Eocene. The 35-degree latitudinal extent of the basin means that it extends across several of Earth’s major climatic zones, from subtropical to subarctic. The basin shifted about 301 northwards between the Mid-Jurassic and the

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Figure 2 Idealized structural and stratigraphic cross-section across the Western Canada foreland basin at a time of maximum transgression. The positions, directions and sizes of arrows indicate relative thrusting, subsidence and uplift (rebound). Modi¢ed from Kau¡man (1977, 1984), with additional terminology from DeCelles and Giles (1996) and Catuneanu et al. (1999, 2000).

Early Cretaceous (Irving et al., 1993) as the North American continent was pushed northwestwards, away from the Mid-Atlantic spreading center. The climate throughout the basin was warmer than at the present day. It steadily cooled during the northward drift, but remained signiﬁcantly warmer than at present (Frakes, 1979). Signiﬁcant latitudinal displacement did not take place after the Early Cretaceous. The Western Interior Basin is characterized by what may be the world’s largest subsurface data base. Petroleum exploration activity commenced here early in the twentieth century, and the number of wells drilled into the basin now numbers in the hundreds of thousands. All the data from the Canadian portion are available in repositories administered by provincial governments. Data access in the United States is less consistent, but is compensated in part by excellent outcrop, especially in the Rocky Mountain states, where strata have been uplifted and exposed as a result of the Laramide fragmentation of the basin (e.g., see Van Wagoner and Bertram, 1995). Intensive study of the basin by American and Canadian geologists has contributed much to the theory of Geology itself, including the following developments: 1. Geodynamic models for foreland basins were ﬁrst developed here, by Beaumont (1981) and Jordan (1981), based on earlier work that was the ﬁrst to recognize the dynamic linkages between crustal loading and basin formation (Price, 1973), and those between the creation of tectonic lands and the development of sediment wedges (Figure 2; e.g., Bally et al., 1966; Price and Mountjoy, 1970). The importance of the ‘‘dynamic load’’ imposed by the mantle was also ﬁrst recognized in this basin (Cross, 1986; Mitrovica et al., 1989). 2. Some of our ideas about large-scale allogenic controls on sedimentation evolved from work carried out here, including concepts of stratigraphic cyclicity (e.g., Kauffman, 1977, 1984) and the importance of regional unconformities (e.g., Weimer, 1960, 1986). 3. Marine invertebrate faunas at the northern end of the basin are typically representatives of the Boreal province, whereas those at the southern end are commonly similar to Tethyan forms. Intermingling of the faunas as climatic belts shifted and sea levels rose and fell has provided invaluable means of correlation between biogeographic provinces, and has made a signiﬁcant contribution to the development of the global chronostratigraphic time scale (e.g., Kennedy and Cobban, 1977). 4. Some of the principles of high-resolution chronostratigraphy have been developed here, including the use of closely spaced datable bentonites to constrain the ages of biozones (Kauffman et al., 1991). 5. The extremely detailed subsurface record, especially that for Alberta (stored at the Energy Resources Conservation Board in Calgary) is permitting some remarkably reﬁned stratigraphic studies, using the principles of high-resolution sequence stratigraphy to explore large-scale allogenic controls on sedimentation (e.g., Plint et al., 1986; Plint, 1990, 1991; Varban and Plint, 2008; Payenberg et al., 2002).

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Figure 3 Precambrian tectonic elements of the Western Interior Seaway basement. Uppercase names and dotted map patterns are Archean provinces, lowercase names are Proterozoic and Phanerozoic orogens. Abbreviations: BH, Black Hills inlier; CH, Cheyenne belt; FNL, Fort Norman line; GL, Great Lakes tectonic zone; GS, Great Slave Lake shear zone; LL, Liard line; MRV, Minnesota River valley;TH,Thompson belt;WR,Winisk River belt. Modi¢ed from Ho¡man (1988, 1989) and Cecile et al. (1997).

Data sources and analyses of the Western Interior Seaway and foreland-basin system are extensive. Several chapters of the GSA Geology of North America series focus on this basin (see especially Aitken, 1993; Peterson, 1988; Molenaar and Rice, 1988), and there have been many special volumes devoted to the geology of part or all of the basin (e.g., Reynolds and Dolly, 1983; Stott and Glass, 1984; Flores and Kaplan, 1985; Macqueen and Leckie,

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Figure 4 Tectonic elements within the Western Interior Basin. These arches and basins were all active at di¡erent times during the history of the seaway, and re£ect reactivation of various basement elements (compiled from numerous sources).

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1982; Caldwell and Kauffman, 1993; Caputo et al., 1994). Much more detailed summaries than there is space for here have been provided by Kauffman and Caldwell (1993): tectonic setting, stratigraphic and paleoceanographic evolution and Beaumont et al. (1993): geodynamic model. DeCelles (2004) provided a detailed discussion of the relationships between magmatism, metamorphism, tectonism and sedimentation in the U.S. portion of the basin. A large-format compilation and atlas of the geology of the Canadian portion of the basin was compiled by Mossop and Shetsen (1994).

2. Geodynamic Framework The Western Interior Basin exists because of the westward drift of the North American plate relative to Europe, and the subduction beneath the western continental margin of the paleo-Paciﬁc Ocean (Panthalassa), which periodically led to shortening and thickening of the cover rocks. The beginning of the deformation that led to the development of the Cordillera is associated with rapid sea-ﬂoor spreading of the Atlantic Ocean, which commenced in the early Middle Jurassic (Monger, 1993; see Chapter 1, this volume). An arc with a backarc basin was established in the California–Nevada area by Early Triassic time (Lawton, 1994; see also Ingersoll, Chapter 11, this volume), and volcanic detritus in strata of Callovian (late Middle Jurassic) age in Utah, Idaho and Montana suggest that magmatic centers had become subaerial and were yielding clastic sediment by this time. However, DeCelles (2004, p. 113) stated that ‘‘the Early to Middle Jurassic retroarc strata do not support the existence of a regional-scale, integrated foreland basin system’’ prior to the Late Jurassic. The ﬁrst-order switch from regional extension to contraction on the western continental margin was modulated, in the Canadian segment, by the successive arrival and collision of a series of terranes and microplates that were carried into the Cordilleran subduction zone between the Mid-Jurassic and the Eocene (Coney et al., 1980; see also Ricketts, Chapter 10, this volume). The welding on (obduction) of successive terranes resulted in extensive underplating of the crust, wedging of crustal slices into older terranes, and delamination of the crust from the autochthonous basement along major, nearly ﬂat-lying de´collement surfaces (Price, 1986; Gabrielse and Yorath, 1991; Cook et al., 1995). Older terranes were displaced eastward with each successive arrival, which renewed the crustal load on the margin of the Canadian Shield and initiated a new cycle of fold-thrust-belt tectonism, uplift, erosion and clastic-wedge generation. Stockmal et al. (1992) compiled a timetable of the terrane collision events and the ages of the major clastic wedges in the Canadian segment of the basin (Figure 5), which indicates a loose correlation between these events, although other processes beyond simple terrane collision also

Figure 5 The six ¢rst-order clastic wedges of the Western Canada foreland basin, shown as a function of time, and the times of accretion of allochthonous terranes. Also shown is the period of development of the Purcell anticlinorium. Adapted from Stockmal et al. (1992).

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occurred, including major transform displacement of the colliding terranes, and changes in the trajectories of the Earth’s major plates, which affected regional intraplate stress patterns. The U.S. portion of the basin was not inﬂuenced in the same way by the successive arrival and docking of offshore terranes (see Ingersoll, Chapter 11, this volume). However, as summarized by DeCelles (2004), between the Late Jurassic and the Mid-Cenozoic, magmatism and tectonism were not continuous in the arc to the west of the basin, and this inﬂuenced the subsidence and sediment-sourcing patterns in the basin. Changes in the angle of the subducting Paciﬁc plate, which inﬂuenced the patterns of igneous activity, were an important inﬂuence, and the impact of basement heterogeneities on the evolution of the thrust belt, are also factors that affected the development of the basin. The Western Interior Basin has the classic asymmetric cross-section of a foreland basin, deepest in the west (W5 km) (where the tectonic load resulted in the greatest deﬂection of the basement ramp), and tapering to the east. The basin is up to 1,500 km wide, extending as far east as Minnesota. This is much wider than is predicted by a simple ﬂexural loading model, as pointed out by Kauffman (1977, 1984). Numerical modeling (Mitrovica and Jarvis, 1985; Mitrovica et al., 1989; Gurnis, 1992) has demonstrated that subsidence in retroarc foreland basins may be ampliﬁed by a ‘‘dynamic, slab-driven’’ effect, resulting from the ‘‘viscous coupling between the base of the continental plate and downward circulating mantle-wedge material that is entrained by the subducting slab’’ (DeCelles and Giles, 1996). This effect can result in considerable widening of the basin, by more than 1,000 km, and additional subsidence of more than 1 km. The rates of dynamic subsidence increase with the rates of subduction, and also increase as the angle of subduction decreases. As these subduction parameters change with time, the rates and wavelength of dynamic subsidence change as well. Catuneanu et al. (1997) referred to this component of subsidence as the ‘‘dynamic load’’ (Figure 2), as distinct from the ‘‘static load’’ of the orogen. Dynamic subsidence was particularly important during the Late Cretaceous, resulting in a signiﬁcant widening of the basin. Compare the isopachs of Early and Late Cretaceous strata in the Rocky Mountain states (Figure 6). Note that the Jurassic section subcrops below the Cretaceous, and that the Cretaceous section extends at least 200–600 km further eastwards onto the craton (Figure 1). Jurassic outliers on the craton east of the map area of Figure 1 are probably cratonic in origin, formed beyond the limit of the back-bulge basin. The distal edge of the axial basin is formed by the forebulge (Figure 2), vertical motion of which is opposite to that of the basin (Beaumont, 1981). The boundary between the basin and the forebulge is a hingeline, which may be located by careful stratigraphic work (Catuneanu et al., 1999, 2000), as noted below. Beyond the forebulge is the shallow back-bulge basin, which is well preserved in some locations (e.g., Morrison Formation of Utah and Colorado; Lower Cretaceous section of Montana–Wyoming–Colorado: DeCelles and Giles, 1996, DeCelles and Currie, 1996). The concept of the forebulge is indicated in Figure 2, but throughout much of the basin, especially the Canadian portion, it is difﬁcult to locate this geodynamic component of the basin. Various tectonic elements

Figure 6 Contrasting styles of subsidence in the foreland of the US Western Interior. (A) The typical pattern of foreland-basin subsidence generated by thin-skinned tectonics and supracrustal loading of the edge of the craton, as shown by isopachs of the Upper Albian--Santonian strata adjacent to the Sevier orogen. (B) Isopach of Campanian--Maastrichtian strata of the same region, showing broad subsidence centered onWyoming and Colorado, attributed to subcrustal loading by a cooling oceanic slab (Cross, 1986).

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Figure 7 The evolving con¢guration of the Western Interior Basin. The original western depositional limit of the basin in the Late Jurassic was uplifted, eroded and cannibalized by subsequent structural deformation, but can still be partially reconstructed from palinspastic reconstruction of thrust shortening, from the evidence of detrital sources and from remnants preserved within thrust sheets.The position of the forebulge at di¡erent stages during basin evolution is shown by the solid colored lines (US data from DeCelles, 2004; Canadian data fromYang and Miall, in press, a, b).

within the Canadian portion of the basin, such as the Sweetgrass Arch (Figure 4) have been suggested as representing part of the forebulge, but none, including this arch, is oriented in the appropriate direction (parallel to the fold-thrust belt), nor shows the appropriate tectonic-stratigraphic history (such as unconformities correlating to times of fold-thrust belt tectonism). The reason for this may be that the forebulge, unlike in the classic basin model (which, ironically, was originally based on this basin; Beaumont, 1981; Jordan, 1981), is commonly buried by the basin ﬁll because of the additional component of dynamic subsidence noted above. Isopach and facies trends have enabled the position of the forebulge to be estimated at different stages of basin evolution, as shown in Figure 7. A conjectural position of this feature during the Late Jurassic (from DeCelles, 2004) places much of the depositional area of the Morrison Formation within the forebulge to back-bulge portions of the basin. Lateral shifts of the forebulge reﬂect the shifting of the supracrustal load and of the magmatic belt to the west, partly in response to changes in the angle of subduction of the Paciﬁc plate. The clastic wedges shown in Figure 2 illustrate schematically a third level of tectonic control of the basin-ﬁll process, that of the relationship between crustal loading and basin subsidence on a regional to local scale. Examples of these clastic wedges are shown in more detail in Figure 8, and their origin is discussed below. The major stratigraphic names used through the basin are shown in Figure 9.

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Figure 8 Stratigraphic cross-section of Cretaceous rocks from central Utah to northeastern Colorado. Thicknesses are based on well and outcrop control.Vertical exaggeration approximately � 151. The Castlegate Sandstone has been interpreted as a product of ‘‘antitectonic’’ sedimentation (Yoshida et al., 1996). Ksx, Sixmile Canyon Formation; Kfv, Funk Valley Formation; Kav, AllenValley Formation; Ksp, Sanpete Formation; Kr, Rollins Sandstone Member; Kcz, Cozzette Sandstone Member; Kco, Corcoran Sandstone Member. From Molenaar and Rice (1988).

Figure 9 Generalized stratigraphic table for the Western Interior Basin, showing only the most well-known of the numerous lithostratigraphic names and the most widespread of the regional unconformities. Derived from numerous sources. Kau¡man’s cycle nomenclature is from Kau¡man (1984) and Kau¡man and Caldwell (1993).

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3. Paleogeographic Evolution A vast body of stratigraphic literature describes the lithostratigraphy, facies, paleogeographic setting and, more recently, the sequence stratigraphy of the basin ﬁll. Many broad trends have emerged from this work, including the persistence of western sediment sources, the spasmodic inﬂuence of basement features on stratigraphic architecture, and the periodic spread of marine waters through much of the basin, as a result of eustatic sea-level rise or regional subsidence. Many researchers have suggested that eustatic sea-level change exerted a major control on basin architecture (e.g., Kauffman, 1984; Schwans, 1995; Van Wagoner, 1995), whereas others (e.g., Krystinik and DeJarnett, 1995) have tested correlations along the length of the basin and could ﬁnd no widespread stratigraphic events that would tend to indicate eustatic control. Regional studies contain many references to such events as the Niobrara, Greenhorn, and other transgressions, and several authors (e.g., Weimer, 1960, 1986; Kauffman, 1984; Leckie and Smith, 1992) have attempted to systematize the regional stratigraphy into systems of cycles that imply regional basin-wide stratigraphic controls. Eustatic sea-level changes are commonly assumed to be the main control. Kauffman’s (1984) system of transgressive–regressive cycles is the best known of these models (Figure 9). In detail, many of these cycles are very complex, including minor transgressive–regressive cycles and local unconformities, and the reality of through-going eustatic events remains to be convincingly demonstrated. Detailed local studies are leading to the conclusion that combinations of tectonic and eustatic mechanisms, on several different time scales, can explain all observed stratigraphic features (e.g., Varban and Plint, 2008). During the Turonian the sea is interpreted to have risen to at least 300 m higher than present levels, generating a paleogeography dramatically different from that of the present (McDonough and Cross, 1991). Undoubtedly there were other periods of high sea-level, as noted below, but it is not yet known whether these reﬂect global eustatic events or were the product of regional tectonism, such as dynamic topographic movements related to mantle thermal processes (see also Burgess, Chapter 2, this volume). To the plethora of stratigraphic names (Figure 9) must be added the nomenclatural unconformity of the Canada–US border, which functions as a surprisingly effective barrier to correlations and a ﬁlter of interpretations of north–south variability within the basin (e.g., see Payenberg et al., 2002). This section attempts to provide a brief summary of the paleogeographic evolution of the basin, based on a series of maps drawn on the same topographic base in order to facilitate comparisons.

3.1. Jurassic In the early Late Jurassic (Oxfordian) a marine seaway occupied much of the future site of the foreland basin system, extending from western Utah through Idaho and Montana to Alberta and Saskatchewan (Figure 10A). Westerly sediment sources associated with contractional tectonism appeared for the ﬁrst time, including the Mesocordilleran Geanticline of Nevada, which was uplifted as a result of the Late Jurassic Nevadan Orogeny (Armstrong and Ward, 1993), and uplifts in Canada associated with the beginning of the Columbian Orogeny (Poulton et al., 1993). Western sediment sources became dominant in the Kimmeridgian (Figure 10B), and this pattern persisted until the ﬁnal uplift of the basin in the Eocene. In southern Alberta, the marine shales and thinly bedded sandstones of the Fernie Formation are overlain by the deltaic, coal-bearing Kootenay Formation. The Fernie represents the early, underﬁlled stage of the basin, and is the only portion of the basin ﬁll that may be assigned to the ‘‘ﬂysch’’ phase of basin development (Stockmal et al., 1992). Progradation of the Kootenay clastic wedge is associated with the arrival and docking of the Intermontane Terrane (Figure 5: Stockmal et al., 1992; Poulton et al., 1993; Ricketts, Chapter 10, this volume), which elevated the Omineca Geanticline (Figure 10B). In the Rocky Mountain states a series of thick ﬂuvial sheet sandstones interbedded with playa-lake deposits and eolian-erg sandstones was deposited, derived from broad, regional uplifts in Idaho and Nevada. These variegated facies are all assigned to the Morrison Formation, a stratigraphic name that is applied across most of the basin south of the US border (Peterson, 1988; Lawton, 1994). DeCelles and Currie (1996) suggested that the forebulge of the basin was located in central Utah at this time, which would indicate that most of the Morrison Formation accumulated in the back-bulge basin. A foredeep may have been present to the west that has now been uplifted and eroded (its speculative position is shown in Figure 7). Conodont alteration colors in pre-foreland basin (Mississippian) strata suggest that between 2 and 5 km of Jurassic strata may have been present prior to Cretaceous thrust-fault uplift and erosion (Royse, 1993; DeCelles, 2004).

3.2. Early Cretaceous Morrison and Kootenay sedimentation continued into the Berriasian, but much of the Berriasian to Barremian (Neocomian) interval is represented by a regional unconformity throughout the Western Interior Basin (Cook

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Figure 10 Paleogeography of theWestern Interior Basin (A) during the Oxfordian and (B) during the Late Kimmeridgian-Volgian.The paleogeography during the Oxfordian indicates the last major in£uence of the craton in determining sedimentation patterns. Sediment sources to the west of the basin are strongly in evidence by the Late Jurassic, and are interpreted as the uplift of regional paleohighs associated with accretion of the Intermontane Terrane.The Mogollon Highland is interpreted as a rift shoulder. Adapted mainly from Poulton et al. (1993) and Lawton (1994). (Note: Same colours are used in Figures 10--12,15,17,19--21.)

and Bally, 1975; Lawton, 1994; Dyman et al., 1994). This period corresponds to a ‘‘magmatic lull’’ in the Cordillera (Armstrong and Ward, 1993), an interpretation recently conﬁrmed by Ducea (2001), who showed that magmatism in the Sierra Nevada reached its lowest level of activity between about 140 and 125 Ma. Stott (1984) interpreted this as a period of ‘‘orogenic quiescence,’’ and McMechan and Thompson (1993) referred to a ‘‘signiﬁcant isostatic readjustment of the craton.’’ Beaumont et al. (1993) explained these events as a normal part

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of the ‘‘post-orogenic exhumation’’ process following terrane collision, during which the orogenic load is reduced by erosion and/or by extensional faulting. Thick Paleozoic cover on parts of the North American craton were removed by uplift and erosion prior to the Late Cretaceous expansion of the Western Interior Seaway, an episode of uplift related to dynamic topographic processes that may have coincided with this period of postorogenic exhumation (Cookenboo et al., 1998; see also Burgess, Chapter 2, this volume). The eastern margin of the Cretaceous basin lies at least 200 km beyond the Jurassic subcrop (Figure 1), indicating the commencement of the widening of the basin by dynamic loading. East of the subcrop, Aptian or Albian strata rest directly on rocks of the Paleozoic craton. In Alberta, Cretaceous rocks overlie Devonian–Carboniferous rocks of the Paleozoic continental margin (the miogeocline). The base of the Cretaceous section, of Late Berriasian or Aptian age, typically consists of a sheet of coarse, ﬂuvial gravels, throughout much of the Western Interior Basin. These deposits are known by various regional lithostratigraphic names (Figure 11), but what they have in common is clear evidence of western tectonic provenance, and eastward paleoﬂow, transverse to tectonic strike, indicating the development of a broad, east-sloping uplift probably not unlike that of Alberta at the present day (Beaumont et al., 1993). The conglomerates rest on a deeply eroded unconformity surface, with local erosional relief of up to 80 m (Cant, 1996). The sheet-like thickness distribution of these conglomerates is not compatible with the wedge-shaped cross-section that would be anticipated from the ﬁll of accommodation generation by supracrustal loading, and it has been suggested that the depositional slope across which the deposits were transported was the result of regional uplift of the magmatic arc (Heller and Paola, 1989). Provenance studies of the foreland-basin strata indicated that following the regional Mid-Cretaceous episode of tectonic quiescence, erosion tapped into oceanic-arc and related rocks, and syndepositional continental margmatic rocks of Quesnellia, far to the west of the orogenic front (Ross et al., 2005). Examination of the ages of these conglomerates, and reconstruction of the subsidence histories suggest that a new phase of ﬂexural loading and subsidence commenced shortly after deposition, initiating a new ‘‘constructive’’ phase of development of the Cordilleran orogen. This included initiation of thrust-fault movement along the Sevier Orogen of Utah–Wyoming–Montana, and is contemporaneous with an increase in the rate of convergence between the North American and Farallon (paleo-Paciﬁc) plates (summarized by Heller et al., 1986, 1988; Lawton, 1994). Above the basal Cretaceous gravels is evidence of the ﬁrst of the series of widespread Mid- to Late Cretaceous marine transgressions that ﬁlled the basin. Isopachs of the Barremian–Aptian interval show the classic wedgeshaped cross-section of a foreland basin foredeep (DeCelles, 2004, Figure 10), with a zone of maximum thickness

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Figure 11 Distribution and paleocurrent directions of Lower Cretaceous gravel in the Western Interior Basin. Correlative Lower Cretaceous units include Ephraim Conglomerate and Cloverly Formation inWyoming, the Lakota Formation in the Black Hills, the Burro Canyon Formation and Lytle Formation in central Utah, Cadomin Formation in Alberta and the Kootenai Formation in Montana. Modi¢ed from a compilation by Heller and Paola (1989).

of more than 1 km centered on central Utah, thinning to less than 50 m over the forebulge (position shown in Figure 7). At least two Cretaceous cycles of transgression occurred in northern Canada (Kauffman and Caldwell, 1993; Dixon, Chapter 16, this volume), but marine waters did not extend southward into the Western Interior Basin until the Aptian (Figure 12). During the Aptian–earliest Albian interval, most of the Western Interior Basin was occupied by ﬂuvial and estuarine systems assigned to such units as the Mannville Group in Alberta–Saskatchewan, and the Kootenai Formation of Montana (note that this unit is not the same as the Lower Cretaceous Kootenay Formation of Alberta). The Mannville Group is characterized by extreme stratigraphic complexity, reﬂecting repeated minor cycles of sea-level change and tectonic adjustment, generating a high-frequency sequence stratigraphy (Cant, 1996; see discussion below). Systems of overlapping and intersecting incised valleys attest to the limited accommodation at this time. The ﬁll of these valleys is typically estuarine or ﬂuvial, and includes numerous stratigraphically isolated sandstone units that constitute the reservoir units of small to medium sized oil ﬁelds. Spectacularly exposed deposits of giant point bars along the Athabasca River in northeast Alberta constitute part of the McMurray Formation, and probably represent the deposits of a major, tidally-inﬂuenced trunk river

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Figure 12 Paleogeography in Aptian--earliest Albian time. Several marine transgressions spread southward from the Arctic during the Early Cretaceous, but open-marine sediments are present only in the northernmost part of the basin, and represent only the latest of these initial cycles. Rivers draining from the Mogollon Highland and the Sevier Orogen, and probably also from the craton, formed an integrated system, which probably included the formation of northward-draining trunk rivers (Lawton, 1994). The Spirit River channel is a reconstruction based on subsurface evidence (McLean and Wall, 1981). The presence of a major river system in northeastern Alberta is indicated by the giant tidally in£uenced point bar deposits of the McMurray Formation (Mossop and Flach, 1983).

draining northwestward from the continental interior (Mossop and Flach, 1983). These deposits constitute the oil sands of the giant Athabasca deposit. A paleogeographic snapshot for this period is shown in Figure 13, and a simpliﬁed cross-section along the axis of the basin is shown in Figure 14. During the period of Mannville sedimentation, the basin was in what Jordan (1995) classiﬁed as an ‘‘overﬁlled’’ state, in that sediment supply outpaced subsidence, the basin remained largely nonmarine, and much detritus was carried by axial drainage systems beyond the main basin depocenter.

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Figure 13 Paleogeographic reconstruction of a typical segment of the proximal margin of the Western Interior Basin during the Aptian, at a time of sea-level lowstand, based on studies of the Mannville Group (Glauconite Formation) in west-central Alberta. Modi¢ed from Rosenthal (1988).

Figure 14 A simpli¢ed cross-section through the Mannville Formation of Alberta--British Columbia, oriented approximately along the Spirit River Channel (the location of which is shown in Figure 12). The Mannville Group constitutes a sequence that spanned about 12--14 Myr. It can be subdivided into higher-order sequences, each of which is represented by a regressive shoreline sandbody and an onlap surface. The Mannville rests with angular unconformity on rocks of Devonian to Jurassic age (from Cant, 1996).

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Kauffman’s ﬁfth marine cycle, the Kiowa-Skull Creek cycle, is represented by a general rise in sea level during the Albian, and led to an interconnection between the cool boreal waters of the north and the warm-temperate to subtropical waters of the south, the transgression extending southward beyond the intracratonic Transcontinental Arch for the ﬁrst time (Kauffman and Caldwell, 1993).

3.3. Late Cretaceous The Upper Cretaceous stratigraphy of the Western Interior Basin is characterized by the deposits of several major marine transgressions. The paleogeographic maps illustrating this period (Figures 15, 17, 19–21) show conditions

Figure 15 Paleogeography of the Western Interior Basin during the Cenomanian. These and subsequent paleogeographic maps have been compiled from numerous sources, including Cook and Bally (1975),Williams and Stelck (1975), Kau¡man (1984), Molenaar and Rice (1988), Leckie and Smith (1992), Lawton (1994) and Mossop and Shetsen (1994).

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Figure 16 A simpli¢ed cross-section through the Dunvegan ‘‘delta.’’ Line of section is shown in Figure 15. The entire Shaftesbury--Kaskapau interval shown in this diagram represents about 3--4 Myr. The major allomember-bounding surfaces in the Dunvegan Alloformation are interpreted as allogenic in origin, whereas the individual (numbered) shingles are interpreted as autogenic products of delta-switching (Bhattacharya and Walker, 1991).

during maximum transgression. During regressive episodes parts of the basin were uplifted and exposed. Gaps in the stratigraphic record are numerous; some represent millions of years, although most are less than one million years in duration (Leckie and Smith, 1992). Eustatic sea-level changes were probably partly responsible for this stratigraphic architecture, but regional and local tectonic processes were also important, as discussed above. Molenaar and Rice (1988) noted that on a 106�7-year time scale the post-Turonian cycles are not synchronous along the length of the Western Interior Basin, and suggested that differences in the timing or magnitude of tectonic events in the basin overprinted any eustatic controls on the regional stratigraphy. Likewise, Krystinik and DeJarnett (1995) could ﬁnd no major correlatable cyclicity in the Campanian–Maastrichtian ﬁll of the basin. Indeed, the concept of regional cyclicity, although ostensibly offering a means to systematize and simplify the stratigraphy of the basin, seems to have been overplayed by some workers. After a brief fall in sea level at the end of the Early Cretaceous, a major transgression occurred during the Cenomanian, forming the very widespread Mowry Shale, followed by the Greenhorn Formation and, in the far eastern limits of the basin, the Ashville Shale. This was one of the most extensive of the great Cretaceous transgressions (Figure 15). Most of the basin was a deep-water muddy sea, but a coastal clastic belt persisted in the west, ﬂanking the Sevier Orogen. Locally, sediment sources of this age formed signiﬁcant clastic wedges, such as the Dunvegan ‘‘delta’’ of northwestern Alberta, consisting of a succession of deltaic wedges bounded by ﬂooding surfaces (Figure 16). During the Turonian the sea reached an all-time high, calculated to be at least 300 m higher than at present (McDonough and Cross, 1991). The paleogeographic reconstruction (Figure 17) shows the sea extending across most of the area now corresponding to the Great Plains, reaching Manitoba and Minnesota and joining with cratonic interior seas occupying the Mississippi Valley and much of the Arctic platform. The map shown in Figure 17 is, in fact, conservative. Shallow seas may have extended over much of the Canadian Shield, with a connection through Hudson Bay to the north Atlantic (Williams and Stelck, 1975; White et al., 2000). The Midto Late Cretaceous was a time of exceptionally high global sea levels, attributed to high rates of sea-ﬂoor spreading (Pitman, 1978) and the rapid dismembering of Pangea, especially along the mid-Atlantic spreading center (Heller et al., 1996). Widespread transgression in the Western Interior Basin led to what has been termed

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Paleogeography of the Western Interior Basin during the Turonian.

the Greenhorn Sea. Extensive carbonate sedimentation occurred in the basin center, extending from the Dakotas south to Nebraska, Colorado, Kansas and Texas (Figure 17). A transect from Utah to Kansas demonstrates that progradational clastic sequences on the basin margin correlate with limestone–marl cycles in the basin center (Figure 18), which has relevance for the nature of allogenic mechanisms in the basin, as discussed below. In Alberta the Cardium Formation was deposited during the Turonian. This thin unit was the subject of one of the ﬁrst studies of high-frequency sequence stratigraphy in the basin (Plint et al., 1986), a study that established the parameters for a major re-evaluation of the basin stratigraphy using sequence-stratigraphic concepts that has been underway since that publication appeared. Figure 19 illustrates a generalized paleogeography during the Coniacian–Santonian, during which the basin was occupied by the Niobrara Sea. The Niobrara Formation, extending from Saskatchewan to Texas, consists

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Figure 18 Turonian sedimentation of the Rocky Mountain states. Correlation of clastic cycles along the basin margin, in Utah, with limestone--shale cycles of the Greenhorn formation in the basin center (Colorado--Kansas). TS, transgressive surface. Line of section is shown in Figure 17. After Elder et al. (1994).

largely of chalky limestone and calcareous shale, passing westward into the thick, monotonous Mancos Shale in Utah and Colorado, and equivalent units along strike to the north. Coarse clastic wedges prograded into the basin from the Sevier Orogen between Montana and Utah, including marginal conglomerate wedges in the most proximal part of the basin (e.g., Indianola Group of Utah: Figure 8). In Utah these conglomerates reach 4.5 km in thickness (DeCelles, 1994; DeCelles and Cavazza, 1999). Regional isopachs for the Coniacian–Santonian interval show ‘‘the beginning of the demise of the well-organized ﬂexural foreland basin system; the foredeep was poorly deﬁned y’’ (DeCelles, 2004, Figure 14 and p. 140). The Campanian was a period of increased tectonic activity along the Sevier Orogen. In Canada this has been attributed to the accretion of the Insular Superterrane, and the uplift of the Purcell Anticlinorium (Stockmal et al., 1992; Ricketts, Chapter 10, this volume). In the Rocky Mountain states, Lawton (1994) argued that increased tectonism occurred as a result of an increase in the rate of orthogonal convergence between the North American and the Farallon plate. The isopach pattern for Campanian–Maastrichtian strata indicates a broad zone of maximum thickness in central Wyoming and central Colorado, far from the Sevier foredeep, suggesting a strong inﬂuence of dynamic subsidence during this time (Cross, 1986; see Figure 6). Flexural loading of the basin created accommodation for the major clastic wedge of the Belly River and Brazeau formations (Figure 20). The northwest end of the basin was ﬁlled by this wedge, and ﬂuvial progradation was to the southeast, along the axis of the basin. A similar pattern of clastic wedge progradation has been mapped in Montana (Two Medicine and Judith River formations) and Utah (Blackhawk Formation and Castlegate Sandstone of the Mesaverde Group: Figures 8, 20). Paleogeographically, the basin was characterized by a broad coastal plain extending along the western, proximal margins from northeastern British Columbia southward at least as far as New Mexico. This pattern persisted into the Maastrichtian (Figure 21), although the Laramide fragmentation of the foreland into separate fault-bounded basins can be detected by stratigraphic thinning, paleocurrent changes and changes in provenance as early as in the Upper Campanian record of Utah (Lawton, 1986; DeCelles, 2004; Lawton, Chapter 12, this volume). The Bearpaw Sea, of Montana and the Canadian Great Plains (Figures 20, 22A), was the last of the great marine transgressions to affect the Western Interior Basin, and was conﬁned mainly to Montana and the Canadian part of the basin. By Maastrichtian time, Laramide movements had fragmented and uplifted most of the US portion of the Western Interior Basin (Dickinson et al., 1988; Ingersoll, Chapter 11; Lawton, Chapter 12, this volume). After the Maastrichtian, global sea levels steadily fell, and the basin itself underwent post-orogenic uplift commencing in the Eocene. The last units to be deposited in the basin are almost entirely nonmarine, including the North Horn Formation of Utah and the Paskapoo and Ravenscrag formations of Alberta–Saskatchewan. A limited marine incursion is represented by the Peace Garden Member in Saskatchewan (Figure 22A). Deposition of the Cypress Hills ﬂuvial gravels continued into the Early Miocene (Leckie, 2006), after which uplift and erosion commenced. Beaumont (1981), citing work on coal moisture and vitrinite reﬂectance, indicated that some 3 km of synorogenic Cenozoic strata have been removed from the Alberta plains by post-Miocene erosion.

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Paleogeography of the Western Interior Basin during the Coniacian--Santonian.

4. Allogenic Mechanisms of Sequence Development Early regional studies of the Western Interior Basin emphasized widespread regional cyclicity and the presence of major regional unconformities (e.g., Weimer, 1960; Kauffman, 1977, 1984). While the importance of

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Paleogeography of the Western Interior Basin during the Late Campanian.

tectonism, speciﬁcally the ﬂexural-loading model, is recognized as the overriding cause of foreland-basin generation, much emphasis has been placed on 106�7-year cycles of eustatic sea-level change as a major controlling mechanism (e.g., Kauffman and Caldwell, 1993; Ryer, 1993; Van Wagoner, 1995; Schwans, 1995). Kauffman’s (1977, 1984) 10 cycles (Figure 9) are the classic expression of this earlier view. However, since the 1980s the global-eustasy model has undergone criticism and re-evaluation (Miall, 1997; Miall and Miall, 2001),

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Figure 21 Paleogeography of the Western Interior Basin during the Maastrichtian.

and speciﬁc studies of basins such as the Western Interior Basin have provided many examples where other mechanisms, particularly regional to local tectonism, seem likely to be important as an allogenic driving mechanism. The growing recognition of the importance of tectonism in generating high-frequency sequences (periodicity o1 million years) is requiring many earlier interpretations to be re-evaluated. At the same time, very detailed local studies are providing increasing evidence for orbital forcing of stratigraphic cyclicity.

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Figure 22 Stratigraphic cross-sections through part of the Upper Cretaceous section of Alberta, showing the reciprocal stratigraphy e¡ect generated by the crustal response to episodes of £exural loading. (A) Cross-section through Campanian to Paleocene strata of Alberta. Note that episodes of continuous sedimentation in the proximal part of the basin correlate with episodes of uplift and erosion in the distal part, over the forebulge, and vice-versa. C, M, P, Campanian, Maastrichtian, Paleocene; e, E, Early; l, L, Late. (B) Details of cross-section across the hingeline of Alberta showing Upper Campanian ammonite zones (Catuneanu et al., 1999, 2000).

4.1. Tectonic cyclicity The review of basin-wide Upper Cretaceous cyclicity by Krystinik and DeJarnett (1995) marked an important departure from most of the earlier literature on the Western Interior Basin. Contrary to the work of Kauffman and others (cited above), which emphasized eustatic controls on sedimentation, Krystinik and DeJarnett (1995) demonstrated that at least in the Campanian to Maastrichtian record, detailed biostratigraphic correlation of seven sections between New Mexico and Alberta does not support the presence of through-going cycles on the 106�7-year time scale. The rate of shortening along the fold-thrust belt and the pattern of resulting subsidence and generation of accommodation for sediment may vary over a 105–107-year time scale for several reasons: (1) Variations due to heterogeneities in the underlying basement. The rate of ﬂexural response to loading is governed largely by the

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strength and elasticity of the underlying crust, and if this varies laterally, subsidence rates will vary as different parts of the crust are loaded during contractional movements, including fold-thrust development and subduction (Waschbusch and Royden, 1992). (2) The kinematics of contractional tectonism. The internal readjustments of crustal blocks as thrusting takes place can cause ﬂexural loading to be episodic and localized (Deramond et al., 1993). However, DeCelles (2004, pp. 150–151) pointed out that ‘‘thrusting events in the Sevier belt were mere increments of the larger deformation ﬁeld that formed the Cordilleran orogenic belt, that regional thrust loading was probably little affected by events in the frontal Sevier belt alone, and that loading was more or less continuous in time rather than sporadic.’’ It is therefore unwise to attempt to correlate individual depositional episodes with speciﬁc local tectonic pulses. Ryer (1993) argued that the areal extent of clastic wedges, their thickness trends, and the nature of associated unconformities can indicate a predominance of eustatic or tectonic control. For example, the Albian Muddy Sandstone of Wyoming and Colorado (Figure 9), and its equivalent in Canada, the Viking Formation, has a thin, sheet-like distribution, contains widespread subaerial unconformities, and is not associated with thick, proximal conglomerate deposits. Ryer (1993) suggested that these features are indicative of a regional drop in sea-level not associated with a major tectonic episode. By contrast, thick clastic wedges, such as the Ferron, and Bluegate Sandstones, with their correlative proximal conglomerate unit, the Indianola Group (Figure 8), are interpreted as tectonically driven. Likewise, Plint and Kreitner (2007) emphasized the signiﬁcance of widespread, thin sequences as indicating eustatic control in their studies of Upper Cretaceous strata in northwestern Alberta. By contrast, Vakarelov et al. (2006) documented a succession containing clear evidence of angular discordance, indicating the importance of structural tilting during sedimentation. Yongtai Yang and Miall (in press, a, b) argued that some widespread ﬁne-grained units, such as the Fish Scales Formation of Alberta, long thought to be a deposit formed during an episode of high sea level, are, in fact, a product of widespread uniform conditions, which developed across the basin during an episode of tectonic quiescence. Clastic wedges have typically been assumed to be ‘‘syntectonic’’ in origin, and their age has been used to infer the timing of major orogenic episodes (see summary and references in Miall, 1981; Rust and Koster, 1984). However, the relationship between tectonism and sedimentation is complex, depending on the balance among a range of controls, including sediment supply and sediment type, and the conﬁguration and rigidity of the ﬂexed basement. Some recent studies of the Western Interior foreland-basin system and elsewhere have demonstrated that tectonism does not necessarily coincide with the progradation of wedges of coarse sediment, but may precede such progradation by a signiﬁcant period of time (Blair and Bilodeau, 1988; Jordan et al., 1988; Heller et al., 1988). In other cases, sediment input and progradation rates are adequate to keep pace with thrust-sheet loading (Sinclair et al., 1991), and may result in the deposition of signiﬁcant, coarse, syntectonic deposits (DeCelles, 2004; Liu et al., 2005). In some cases, however, the immediate basinal response to crustal loading may be marine or lacustrine incursions, with ponding of coarse debris against the basin margin. Tectonism eventually leads to increased basin-margin relief and hence to a rejuvenation of the supply of coarse sediment, but this process takes time, and in such cases, clastic-wedge progradation is largely a post-tectonic phenomenon (Blair and Bilodeau, 1988; Heller et al., 1988). Areally extensive coarse ﬂuvial deposits may not be deposited until posttectonic uplift of the basin takes place, driven by isostatic rebound following erosional unrooﬁng of the foldthrust belt. Heller and Paola (1992) referred to this model of coarse sedimentation as ‘‘antitectonic.’’ There has been considerable debate in the literature regarding the validity of the syntectonic and antitectonic models with respect to particular foreland basins (e.g., Lageson and Schmitt, 1994; see also summary in Miall, 1996, Section 11.3.6). Varban and Plint (2008) noted a contrast between the wedge-shaped allomembers of the Cenomanian– Turonian Kaskapau Formation of northwestern Alberta and northeastern British Columbia, and the more tabular shapes of the overlying Cardium allomembers. The overall architecture of the Kaskapau allomembers suggests accumulation in a subsiding foredeep under the inﬂuence of crustal loading, the wedging out of individual units away from the orogenic front reﬂecting thinning and onlap of the forebulge. By contrast, the overlying Cardium Formation is interpreted to have been deposited at a time of greater tectonic quiescence. During Cardium sedimentation, sheets of gravel were spread for hundreds of kilometers into the basin from the orogenic front by river systems, to be reworked by marine processes during transgressive episodes, whereas coarse clastics in the underlying Kaskapau are conﬁned to the proximal, western part of the basin. This contrast in stratigraphic architecture and the distribution of coarse clastic is consistent with the two-phase model of foreland basin development suggested by Heller et al. (1988). The reconstructed cross-section shown in Figure 8 illustrates the Upper Cretaceous clastic wedges of the Rocky Mountains states, extending from the Gunnison Plateau, south of Salt Lake City, eastward to the edge of the Great Plains, east of Denver. The section encompasses about 25 Myr, and illustrates the general eastward progradation of coarse clastic facies in response to the progressive shortening of the Sevier Orogen. In detail, progradation took place in a series of shorter pulses, each representing about a million years, or less. Most or all of

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these progradational pulses were driven by local to regional episodes of ﬂexural loading and rejuvenation of sediment sources and clastic delivery pathways (DeCelles and Giles, 1996; DeCelles et al., 1995), but it remains unclear in each case whether these were syntectonic or antitectonic in origin. Yoshida et al. (1996), in a detailed examination of the Castlegate Sandstone, of Campanian age, and building on the tectonic-accommodation model of Posamentier and Allen (1993), argued that its widespread, sheet-like architecture suggests an antitectonic origin. Likewise, the widespread gravel units that formed throughout the basin in the Early Cretaceous (Figure 11) are interpreted as the product of a major episode of uplift and erosion of the orogen predating a new round of ﬂexural subsidence (Beaumont et al., 1993). A detailed study of the Upper Cretaceous proximal foredeep deposits in western Wyoming enabled Liu et al. (2005) to recognize ﬁve megasequences averaging 6.2 Myr in duration, the composition and architecture of which could be related to major thrusting episodes along the Sevier orogen. Where large data bases are available (such as the subsurface data for Alberta, and the excellent outcrop around the San Juan basin of New Mexico and the Book Cliffs of Utah: Figure 8), it is possible to deﬁne and map a highfrequency sequence stratigraphy indicating an episodicity or cyclicity of allogenic processes over time scales of less than one million years. This permits an extension of Ryer’s (1993) ideas about interpreting sedimentary controls from stratigraphic architecture to a much ﬁner scale. Units with a blanket architecture in an area of known tectonic activity (a foreland basin is such an environment, by deﬁnition) suggest a uniform creation of accommodation, a process most likely associated either with eustatic rise in sea level, or regional, epeirogenic, subsidence generated by dynamic topography. Linear isopach highs in ﬁne-grained facies, outboard from a coastal-plain environment, suggest clinoform progradation. Unusual thickness trends and anomalous thickness changes may be associated with active basement tectonism. Wedges thickening toward the fold-thrust belt indicate orogenic sourcing (the ‘‘syntectonic’’ model). The Mannville Group (Figure 14) exhibits the architecture of at least two, possibly three scales of sequence cyclicity (Cant, 1996). The overall succession is that of a transgressive to highstand systems tract developed as a result of a cycle of accommodation change lasting a few million years. Nested within this are high-frequency units, as shown in Figure 14. Cant (1996) interpreted the latter as resulting from cycles of less than one-million years duration. Where sediment supply was low, the succession consists largely of superimposed incised-valley ﬁlls. The Dunvegan delta (Figure 16) was generated by a similar cyclicity, resulting in a series of widespread allomembers (members A to G of the Dunvegan Formation in Figure 16), which have been traced for several hundred kilometers. Plint (1991) demonstrated a similar high-frequency sequence stratigraphy in the units above the Dunvegan Formation. He tentatively interpreted the allogenic mechanism as glacioeustasy. Episodes of ﬂexural loading increase the amplitude of the crustal deﬂection, which results in uplift of the forebulge. Cessation of loading, as at the end of an episode of thrust movement, causes relaxation, and subsidence of the forebulge. Catuneanu et al. (1999, 2000) demonstrated that these processes are recorded by facies changes in the stratigraphic record. Periods of active ﬂexural subsidence in the foredeep are represented by stratigraphic successions that ﬁne upward, indicating increasing accommodation, whereas decreasing accommodation over the rising forebulge is indicated by upward coarsening. The opposite facies trends occur during episodes of ﬂexural relaxation. Using bentonite horizons as markers, Catuneanu et al. (1999, 2000) traced individual cyclic units that showed a reversal in vertical facies trends where they crossed the hingeline at the edge of the forebulge (Figure 2). These sequences indicate alternating episodes of subsidence and uplift of the basin over periods ranging from somewhat more than one million years (Figure 21A) to somewhat less than one million years (Figure 9.22B). This is a particularly useful demonstration conﬁrming the reality of high-frequency tectonism related to the ﬂexuralloading process. Sea-level rise and fall could not generate vertical facies changes that reverse in trend where a sequence crosses from the basin to the forebulge. Crustal shortening may generate folding and faulting within the basin ﬁll and, as Zaleha et al. (2001) demonstrated, based on their studies of the Lower Cretaceous deposits of Wyoming, movement on syndepositional structures may have measurable effects on ﬂuvial channel trends and channel architectures. Changes in the intraplate stress ﬁeld emanating from the fold-thrust belt were also likely the cause of subtle syndepositional changes in stratigraphic thicknesses and paleoslopes (Heller et al., 1993). Regional mapping demonstrates that, in plan view, the hingeline at the edge of the forebulge has an arcuate trend, the point at the center of the arc corresponding to the main locus of loading during each thrusting episode (Figure 23). This locus moved northward several hundred kilometers between the Campanian and the Paleocene, which is consistent with the right-lateral transcurrent deformation that the orogen was undergoing at this time (Monger, 1993).

4.2. Basement control A recurring theme in discussions of the Western Interior Basin (e.g., Ross and Eaton, 1999; Gay, 2001) has been controversy regarding the extent to which thickness and facies trends in the Phanerozoic cover parallel terrane

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Figure 23 Location of the hingeline between the foredeep and the forebulge during the Campanian--Paleocene. The arcuate trend of the hingeline indicates the locus of greatest £exural load in the orogen---at the center of the arc. C, M, P, Campanian, Maastrichtian, Paleocene; e, E, Early; l, L, Late.The location of maximum loading shifted progressively northward during the Late Cretaceous--Paleocene (Catuneanu et al., 1999, 2000).

boundaries and other ancient structures in the basement (Figure 3). Locally the evidence commonly is equivocal, and needs to be examined on a case-by-case basis. Pang and Nummedal (1995) demonstrated the importance of local ‘‘buttresses’’ and ‘‘zones of weakness’’ in affecting ﬂexural subsidence patterns and, as Heller et al. (1993) have suggested, ‘‘changes in intraplate stresses add a small, but stratigraphically signiﬁcant, component of uplift or subsidence to preexisting topography and/or zones of weakness within the lithosphere.’’ Yoshida et al. (1996) noted the possible inﬂuence of heterogeneity related to the underlying Paradox Basin as a cause of differential movement during the Late Cretaceous, and the Douglas Creek Arch, Uncompahgre Uplift (Figure 4), and other pre-Cretaceous structural features (see Blakey, Chapter 7, this volume) are all elements that potentially could have subtly inﬂuenced sedimentation, including paleocurrent trends and thickness patterns. At the local scale, paleogeographic patterns, isopach and isolith trends, and paleocurrent dispersal trends may reﬂect basement control, and may be important guides to interpretation. For example, Hart and Plint (1993) and

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Plint et al. (1993) discussed tectonic mechanisms that governed construction of the Smoky Group in the foreland basin of Alberta and British Columbia. The overall architecture of the Smoky Group was interpreted as a product of a varying rate of ﬂexural subsidence and associated forebulge movement on a 106-year time scale (Plint et al., 1993). Bevels on certain erosion surfaces within the Cardium Formation were interpreted as the result of ﬂexure over reactivated basement faults and resulting shoreline incision (Hart and Plint, 1993). Donaldson et al. (1998) interpreted subtle drape effects within Cretaceous strata of NW Alberta as the result of reactivation by forelandbasin tectonism of faults cutting the Precambrian basement.

4.3. Milankovitch cyclicity There is increasing evidence for basin-wide cyclicity forced by Milankovitch processes, particularly in the Cretaceous record of the Western Interior Basin. Orbitally forced climate change affects rainfall and vegetation cover, and hence sediment yield. It can also affect oceanic circulation patterns, redox potential and organic productivity (Fischer, 1986). The evidence for orbitally forced eustatic sea-level change remains controversial, with most recent workers (e.g., Markwick and Rowley, 1998) arguing that there is no evidence for signiﬁcant land ice during the Mesozoic. However, the presence of a high-frequency cyclicity (episodicity of less than one million years) that can be traced for long distances, and correlated across different climatic and/or tectonic zones, is suggestive of Milankovitch control, especially if it occurs in basincenter sites where high-frequency tectonically-controlled variations in sediment supply can be ruled out as improbable. One of the best examples of probable Milankovitch cyclicity within the Western Interior Basin is the succession of seven Turonian sequences mapped by Elder et al. (1994) between Utah and Kansas (Figure 18). At the western margin of the basin, in Utah, a series of coarsening-upward successions consist of marine mudstone, siltstone and silty sandstone, passing up into crossbedded shoreline sandstone. Pebbly sandstone lag deposits form the top of each of these cycles. Based on biostratigraphic data and correlation of bentonite markers, Elder et al. (1994) correlated these successions with limestone-marl couplets in Kansas, and suggested correspondence to an obliquity or eccentricity periodicity. Their interpretation is that cool and/or wet phases of an orbital cycle would have increased clastic sediment yields from the orogen, leading to increased rates of coastal progradation and increased clay sedimentation in the basin center, whereas warm and/or dry phases would have reduced sediment yield, and would have favoured carbonate production in the basin center. Tectonic mechanisms for this cyclicity seem unlikely, given the wide distribution of these cycles and the fact that they can be traced through major facies changes. However, Ulicˇny (1999) questioned some of the correlations suggested by Elder et al. (1994) and emphasized the possible importance of tectonism as a driving mechanism for these sequences. The role of eustasy in the generation of these cycles is unclear. Small ﬂuctuations in ocean temperatures between cool and warm phases of the cycles may have led to volume changes in the oceans, which may have been enough to cause small changes in eustatic sea levels. A detailed, quantitative geochemical analysis of limestone–marlstone couplets in the Bridge Creek Limestone Member of the Greenhorn Formation was performed on a core drilled in central Colorado by Sageman et al. (1997). The data were amenable to spectral analysis, which demonstrated the presence of all three of the main Milankovitch frequencies in the succession, and conﬁrmed the importance of orbital forcing as a mechanism for generating high-frequency sequences in the basin. The highly detailed sequence-stratigraphic studies being carried out by A. G. Plint and his research team provide the basis for detailed analyses of the question of cyclic mechanisms. In a recent study, Varban and Plint (2008), as noted above, provided a convincing demonstration of the importance of ﬂexural subsidence and the inﬂuence of a forebulge during the deposition of an Upper Cretaceous succession spanning about 5.6 Myr. Accommodation generation kept approximate pace with subsidence, so that the shallow-marine sedimentary environments did not change much through this period. Detailed stratigraphic analysis led to the subdivision of the section into some thirty-eight allomembers that could be traced for some 400 km, across the entire foredeep. Many of the allomember boundaries are best interpreted as the product of forced regression, and are succeeded by facies indicating ﬂooding and backstepping (Plint and Kreitner, 2007). Cycles of eustatic sea-level change with frequencies varying between about 10 and 125 ka seem to be indicated. Given the extremely low depositional slopes, an amplitude of 10 m for the sea-level cycles was estimated by Varban and Plint (2008) to have been sufﬁcient to generate the observed stratigraphic architecture. They referred to other studies of Milankovitch cyclicity during the Cretaceous, which are contributing to an increasingly convincing case for the growth and decay of small continental ice caps during what has for some time been assumed to have been a long-term phase of global greenhouse climatic conditions.

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4.4. Discussion Current research trends do not support the concept of basin wide 106�7-year cyclicity (what has been called second- to third-order eustasy) in the Western Interior Seaway. Stratigraphic cyclicity is pronounced, at this time scale, but is local to regional in extent, and is a product of tectonism, including contractional movement of major thrust sheets, foredeep subsidence, forebulge uplift, and migration of these tectonic elements as a result of crustal shortening. Evidence for mild warping and tilting as a result of intraplate stresses can also be demonstrated. Documentation of high-frequency sequences with periodicities of less than 1 million years in many parts of the Western Interior Basin (e.g., see discussion by Bergman and Walker, 1995, p. 263) has renewed the debate regarding allogenic mechanisms. There is now excellent evidence for both high-frequency tectonism (Catuneanu et al., 1999, 2000; Vakarelov et al., 2006) and orbital forcing (Elder et al., 1994; Sageman et al., 1997; Plint and Kreitner, 2007) as driving mechanisms at different times, in different parts of the basin. Future work will need to focus on very reﬁned correlations of sequences between tectonically and paleogeographically contrasting parts of the basin, as a basis for attempts to resolve the question of causality in each case.

5. Economic Resources 5.1. Oil and gas Approximately 40% of oil and gas production from the Rocky Mountain states is from Cretaceous reservoirs, mainly shallow-marine (shelf and shoreline) sandstones (Molenaar and Rice, 1988). Source beds are present in such units as the Mowry Shale, Greenhorn Limestone and Niobrara Formation. Major oil ﬁelds with reserves in excess of 16 � 106 m3 (100 million barrels) occur in Upper Cretaceous reservoirs in Wyoming and Montana. The second largest gas ﬁeld in the conterminous 48 US states is the Blanco gas ﬁeld of the San Juan Basin, in northwestern New Mexico. Ultimate production from this ﬁeld is estimated at 650 � 109 m3 (23 tcf ) of gas, producing primarily from the ﬂuvial-deltaic Dakota Sandstone and the coastal barrier sandstones of the Mesaverde Group. The Western Canada foreland basin system contains an estimated 1 � 109 m3 (6.4 � 109 barrels) of recoverable reserves of conventional oil and gas (Porter, 1992). Cretaceous reservoirs contain 26.6% of the oil in the basin, and 48.1% of the marketable gas reserves (Hay, 1994). Mannville Group shallow-marine and ﬂuvial–estuarine sandstones host most of the ﬁelds, most of which are relatively small (o13 � 106 m3 or o80 million barrels of oil, or gas equivalent). The deep-basin Elmworth gas ﬁeld, in west-central Alberta, is the largest gas ﬁeld (31 � 109 m3 or 1092 bcf) within this group. Signiﬁcant reserves are also hosted in the Viking and Cardium sandstones, the Pembina ﬁeld, a Cardium ﬁeld, near Edmonton, being the largest (238 � 106 m3 or 1.5 � 109 barrels). Important gas reserves are present in Upper Cretaceous shallow-marine sandstones in southern Alberta. The Canadian oil sands reserves are amongst the largest of their kind in the world. The three major deposits in north-central Alberta are all hosted in Mannville Group (Lower Cretaceous) rocks and are estimated at 207 � 109 m3 (1.3 � 1012 barrels) of oil and bitumen in place (Porter, 1992).

5.2. Coal Upper Cretaceous rocks of the Western Interior Basin contain large reserves of coal, which are under active exploitation for thermal power generation. Much of the coal is obtained by strip mining. Coals in Utah and Colorado are largely of medium- to high-volatile bituminous rank, those of Arizona and New Mexico are of sub bituminous rank (Molenaar and Rice, 1988). In Canada the major coal-producing units are the Jurassic Kootenay Formation, the Lower Cretaceous Gething Formation and Fort St. John Group, and the Upper Cretaceous Brazeau Group in the Foothills (Smith et al., 1994).

6. Conclusions This chapter has attempted to provide a broad, panoramic view of the world’s most well documented large foreland basin, the study of which has contributed much to an understanding of the relationships between plate– tectonic processes, regional tectonism, magmatism and sedimentation on the western margin of North America as well as to the advancement of geology itself (as noted in the Introduction). While it is cliche´ that stratigraphic successions store the record of long past geological processes, the establishment of the temporal relationships

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between any geological process and a speciﬁc sedimentary product remains both a signiﬁcant challenge and one of the most useful tests of many geological hypotheses. As with so many other questions in Geology, such tests point to the need for ever more precise chronostratigraphic documentation. The debate regarding allogenic controls on sedimentation is particularly well illustrated by the many discussions that have taken place regarding the origins of speciﬁc sedimentary units in the Western Interior Basin, because of the detailed mapping that has been carried out here. But much remains to be done. The re-evaluation of basin stratigraphy using sequence-stratigraphic principles, initiated by A. G. Plint in the 1980s, is far from complete. Regional crustal studies, such as those carried out by Canada’s Lithoprobe project, and by COCORP in the United States, also remain to be exploited for their full implications. The huge subsurface data base, particularly that available for the Canadian portion of the basin, and the superb outcrop in the Rocky Mountain states, where the foreland basin has been disrupted, uplifted and exposed by the Laramide movements, offer opportunities for fruitful research for many years to come.

ACKNOWLEDGMENTS We acknowledge the useful comments provided by the critical readers, Ron Blakey, Ray Ingersoll and Peter DeCelles.

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D., 1983, Deep channel sedimentation in the Lower Cretaceous McMurray formation, Athabasca oil sands, Alberta. Sedimentology, v. 30, pp. 493–509. Mossop, G. D., and Shetsen, I. (compilers), 1994, Geological Atlas of the western Canada Sedimentary Basin, Canadian Society of Petroleum Geologists, Calgary, AB, 510 pp. Pang, M., and Nummedal, D., 1995, Flexural subsidence and basement tectonics of the Cretaceous Western Interior Basin, United States. Geology, v. 23, pp. 173–176. Payenberg, T. H. D., Braman, D. R., Davis, D. W., and Miall, A. D., 2002, Litho- and chronostratigraphic relationships of the Santonian/ Campanian Milk River Formation in southern Alberta and Eagle Formation in Montana utilizing stratigraphy, U-Pb geochronology and palynology. Canadian Journal of Earth Sciences, v. 39, pp. 1553–1577. Peterson, F., 1988, A synthesis of the Jurassic system in the southern Rocky Mountains, in Sloss, L. 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CHAPTER 10

Cordilleran Sedimentary Basins of Western Canada Record 180 Million Years of Terrane Accretion Brian D. Ricketts

Contents 1. Introduction 2. The Cordilleran Morphogeological Belts 3. Terranes, Terrane Accretion and Associated Basins of the Canadian Cordillera 3.1. Intermontane superterrane 3.2. Insular superterrane 3.3. Magmatism, deformation, and relative plate motions 3.4. The modern plate boundary 3.5. The ‘‘Baja BC’’ debate 4. Sedimentary Basins Associated with Intermontane Superterrane 4.1. Whitehorse trough 4.2. Bowser Basin 4.3. Sustut ‘‘piggyback’’ Basin 4.4. Tyaughton–Methow basin 5. Basins Located along the Inboard Margin of Insular Superterrane 5.1. Nutzotin–Dezadeash–Gravina–Gambier basins 6. Basins Located along the Outboard Margin of Insular Superterrane 6.1. Queen Charlotte–Wrangell Mountains basins 7. Cenozoic Basins-Harbingers of the Modern Plate Boundary 7.1. Queen Charlotte-Georgia-Toﬁno basins 7.2. Tertiary Queen Charlotte Basin 7.3. Georgia (Nanaimo) Basin 7.4. Toﬁno Basin 7.5. Provenance linkages 8. Discussion Acknowledgments References

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Abstract The North American Cordillera is a collage of disparate and commonly far-traveled lithospheric blocks and slivers (terranes). Beginning in the early part of the Jurassic, terranes accreted to the western margin of the North American craton such that the margin has grown several hundred kilometres westward to its present position. Terrane accretion continues today. Jurassic to Recent foreland, forearc, backarc, wrench, and remnant-ocean basins in the Canadian Cordillera, west of the Foreland Belt, record complex relationships between terrane accretion to ancestral North America, and the crustlithosphere responses (subsidence, uplift, denudation) associated with collision, subduction, rifting, and wrench tectonics. Basin subsidence, driven for example by tectonic loading during terrane collision, may be terminated by accretion of successive terranes and terrane-amalgams (superterranes). The signiﬁcance of successive accretion events for basin evolution is well illustrated in Bowser Basin where subsidence and sedimentation were associated with the interaction among Stikinia, Quesnellia, and Cache Creek terrane (components of the Intermontane superterrane) and ancestral North America, beginning in the early Middle Jurassic. Crustal shortening across Bowser Basin beginning in the Early Cretaceous, was likely driven by the docking farther west of Insular superterrane. Indeed, it is likely that tectonic Sedimentary Basins of the World, Volume 5 ISSN 1874-5997, DOI 10.1016/S1874-5997(08)00010-5

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loading and accumulation of major clastic wedges in the Western Canada foreland basin in Alberta and British Columbia also were mechanically linked to Intermontane and Insular superterrane accretion. It is generally accepted that most Cordilleran terranes traveled some distance prior to accretion. Some, like Stikinia, essentially traveled alone; others, like the Insular superterrane contained lithospheric blocks that were amalgamated before docking with North America. However, vigorous debate continues concerning the distances and paleolatitudes traversed by each terrane, and to some extent the timing of terrane accretion to the North American plate margin.

1. Introduction Sedimentary rocks making up the mountainous swath that is the Canadian Cordillera constitute basins or remnants of basins dating back to the Precambrian. However, it is only since the beginning of the Middle, and possibly Early Jurassic that any of these Cordilleran basins has formed in response to plate-tectonic processes acting directly on the ancient western margin of the North America. All Cordilleran basins older than this formed on plates remote from the North American margin and subsequently were accreted to it. The North American Cordillera is a collage of disparate and commonly far-traveled lithospheric blocks and slivers (terranes) that have accreted to the western craton margin such that the margin has grown several hundred kilometres westward to its present position (Figure 1). A terrane is deﬁned as a fault-bounded block containing rocks that have a distinct geologic history compared with contiguous blocks. Howell (1995) deﬁned different kinds of terranes as: stratigraphic representing fragments of continents, continental margins, volcanic arcs, and oceanic crust; disrupted characterized by pervasive shearing and penetrative deformation; and metamorphic. Terms like exotic and

Figure 1 (a) Map of the principal terranes discussed in this chapter, the accretion of which over the past 180--200 Myr has given rise to the Canadian Cordillera (adapted from Jones et al., 1986; Monger, 1989). Terrane abbreviations are listed in Table 2. (b) The ¢ve morphotectonic belts making up the Canadian Cordillera. The western limit is the modern plate boundary to North America; the eastern limit is the stable North American craton. The main characteristics of each belt are summarized in Table 1.

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suspect generally are used for terranes of unknown origin and, with the term allochthonous, imply tectonic transport. In contrast, pericratonic terranes contain rocks that are probably related to an autochthonous craton but may overlie attenuated crust and have uncertain paleogeographic afﬁnities. Additional terms that are important to the Cordilleran story are composite or superterranes, which apply to two or more terranes that have amalgamated into a single lithospheric entity before ﬁnal accretion to the North American margin. Two superterranes, Intermontane and Insular superterranes, contain most of the basins in the Cordilleran collage of western Canada. Plate-tectonic processes involving Cordilleran terranes were predominantly contractional, including subduction and obduction, collision, crustal-scale delamination, and transcurrent or strike-slip faulting. A sense of the magnitude of lithospheric convergence along the western margin of North American is illustrated by Engebretson et al. (1992) who calculated that about 13,000 line-km of mostly oceanic crust have been consumed in subduction zones over the last 200 Myr. The Jurassic and younger Cordilleran basins provide a record of these terrane-accretion processes in the following interrelated ways: the timing of basin-subsidence records crustal loading by contraction or extension; basins or stratigraphic assemblages that extend across adjacent terrane boundaries (also called overlap or successor basins) help to bracket the timing of the accretion events; sedimentation style and stratigraphic architecture provide critical information on paleogeography and the capacity for the basin to accommodate sediment; sediment composition provides a record of source rock and changes in source resulting from tectonic relocation or unrooﬁng of terranes; paleontological and isotopic-radiometric records provide temporal constraints for accretion events; fossils also provide a measure of control on terrane paleolatitude prior to and during accretion. Paleomagnetic data provide independent constraints on terrane paleolatitudes. The remainder of this chapter is organized as follows: elucidation of the ﬁve Cordilleran morphogeological belts is followed by a conspectus of the major Cordilleran terranes, their accretion to North America, and the resulting sedimentary basins. The major basins are then examined in relation to associated terrane-accretion events in terms of their subsidence history, stratigraphic and sedimentologic architecture, and provenance (Figure 2).

2. The Cordilleran Morphogeological Belts The Canadian segment of the Cordillera has been divided into ﬁve belts. Each belt is identiﬁed by the congruence between its distinctive bedrock geology and its physiography; each belt is oriented parallel to the overall tectonic trend (Figure 1) (Wheeler and Gabrielse, 1972; Monger et al., 1972; Tipper et al., 1981). As plate-tectonic theory and subsequent models of terranes developed, it was recognized that the extent of the morphogeological belts correspond closely to the boundaries of two superterranes (the Insular and Intermontane superterranes), the two intervening swaths of plutonic and metamorphic rocks (the Omineca belt and Coast belt; Monger et al., 1982), and the easternmost Foreland belt that tapers eastward into undeformed rocks that overlie the stable craton (Figure 1). The main characteristics of each morphogeological belt are summarized in Table 1. Over much of their length the Coast and Omineca belts are bounded by fold-and-thrust belts. Folding and thrusting in the Foreland belt, including subsidence of the Alberta Foreland basin, is dynamically and mechanically linked to the Omineca belt (Price, 1973; Stockmal et al., 1992; Miall et al., Chapter 9, this volume). Likewise, the Coast Belt Thrust System is thought to be dynamically linked to accretion of the Insular Superterrane (Crawford et al., 1987; Journeay and Friedman, 1993). The Intermontane and Insular belts are underlain by large areas of volcanic (commonly arc and forearc assemblages) and sedimentary strata, most of which have been metamorphosed to sub-greenschist grades, in contrast to rocks in the adjacent ‘‘crystalline’’ belts. The sedimentary basins examined herein are located in or at the margins of the Intermontane and Insular belts (Figure 2). Development of each basin was linked to terraneaccretion events and associated crustal-scale contraction or extension, lithospheric ﬂexure, and arc magmatism.

3. Terranes, Terrane Accretion and Associated Basins of the Canadian Cordillera The North American Cordillera is made up of more than 200 terranes (Jones et al., 1986; Monger 1989), ranging in size from a few kilometres to the length of British Columbia (Stikinia). The terranes themselves consist of one or more tectonic assemblages, bounded by faults or unconformities, in which strata have a common depositional setting (e.g., platform, shelf, volcanic arc or forearc) and/or a common structural fabric or tectonic history (Tipper et al., 1981; DNAG map 1712A). The deﬁning characteristics of terranes relevant to this synopsis are shown in Table 2. An encyclopaedic compilation of Canadian Cordilleran geology is contained in the Decade of North American Geology Cordillera

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Figure 2 An outline of the sedimentary basins discussed in this chapter (the map underlay is from Figure 1a. Adapted from Eisbacher (1985), Monger (1989), McClelland et al. (1992). Terrane abbreviations are listed in Table 2.

Orogen volume and accompanying maps (Gabrielse and Yorath, 1991), in particular Terrane Map 1713A (DNAG). The timing of terrane–superterrane accretion and associated basin formation are summarized in Figure 3. Provenance linkages that correspond with these events are shown in Figure 4.

3.1. Intermontane superterrane Intermontane and Insular composite (super)terranes, their amalgamation, and eventual accretion to North America are central to understanding the history of Canadian Cordilleran basins. However, uncertainties and

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Table 1

Deﬁning characteristics of the ﬁve morphogeological belts of the Canadian Cordillera (see also Figure 1b).

Insular Belt West

Coast Belt

Inter-Montane Belt

Omineca Belt

Foreland Belt East

Western boundary at Mountainous terrain Mostly low relief, High relief uplifts of Middle Proterozoic to the base of the underlain by except in the Skeena plutonic and highPaleogene modern continental extensive plutonic & area of north British grade metamorphic sedimentary rocks, slope & the Queen metamorphic rocks. Columbia. E and W rocks, and slivers of overlying North Charlotte transform Plutons tend to boundaries coincide sedimentary strata America crystalline fault; the eastern become younger to with major changes in having afﬁnities with basement deformed boundary extends the east. Mostly topographic relief the Foreland Belt. into an E-NE verging from east Vancouver I quartz diorite, diorite Sedimentary strata are Pervasive penetrative thrust stack. Up to to the Denali Fault & & tonalite; Sr initial mostly sub-greendeformation. Core 200 km shortening. Saint Elias Mts. in the ratios o/ ¼ 0.704. schist grade. zones of Western boundary is north. Generally, low Metamorphic grade Deformation style metamorphic and the N and S Rocky grades of ranges from greendepends on basement rock. Mountain Trenches metamorphism schist to amphibolite competency of rock and eastern Selwyn units; e.g., tight- to Mts. in the north. open-folds in the Sedimentation and Skeena Fold Belt deformation are linked to terrane accretion events Source: Information is from Monger et al. (1972, 1982) and Gabrielse et al. (1991)

debate remain over the timing of these events and the horizontal distances terranes may have traveled from some paleolatitude before accretion took place. Slide Mountain terrane, the easternmost component of Intermontane superterrane (Figure 1a), contains Upper Paleozoic and some Triassic volcanic, ultramaﬁc, and sedimentary rocks that have ocean-basin afﬁnities. Initial thrusting over Omineca belt and the pericratonic Kootenay and Cassiar terranes was probably Middle Jurassic in southern British Columbia, but may have been earlier farther north. Rare clasts of metamorphic rock having pericratonic afﬁnities suggest close links between the Slide Mountain ocean and the adjacent craton margin. However, contradictory evidence from Permian faunas suggests that original paleolatitudes may have been as far south as northern Mexico (Monger et al., 1991). Slide Mountain terrane was separated from Cache Creek oceanic crust by Quesnellia (Figure 1a). Quesnellia volcanic and volcaniclastic assemblages have island-arc afﬁnities spanning the Late Triassic through Middle Jurassic. Low strontium isotope initial ratios (o0.704) indicate a generally primitive magmatic character (Armstrong, 1988). Cache Creek terrane, located west of Quesnellia, is similar to Slide Mountain terrane in that it comprises igneous and sedimentary lithologies having ocean-basin afﬁnities. In central British Columbia, Cache Creek ocean-basin strata have been deformed into a crustal-scale thrust stack that bears the hallmarks of an accretionary prism above a subduction complex (Struik et al., 2001). The complex forms part of the Pinchi suture which Struik et al. (2001) interpreted as the lithospheric-scale collisional boundary. Stikinia, the largest terrane in the Canadian Cordillera, is located west of Cache Creek terrane. Like Quesnellia, it consists predominantly of rocks having volcanic arc afﬁnity. In northern British Columbia, Stikinia is overthrust by the Cache Creek terrane along the King Salmon thrust. The timing of Slide Mountain-Quesnellia-Cache Creek-Stikine terrane amalgamation (Figure 3) and the timing of accretion of the resulting Intermontane superterrane to the western margin of North America is still debated. Most agree that there was some interaction among the terranes as early as Late Triassic (e.g., Monger, 1989). Whitehorse Trough in the northern Intermontane belt (Figure 2), is an Early Jurassic arc-marginal or forearc basin that overlies Stikinia and Cache Creek terranes, is located northeast of an active Stikinian arc, and provides evidence of arc-subduction links as early as Sinemurian among Cache Creek, Stikinia, Quesnellia, and the ancient North American margin (Figure 3) (Tempelman-Kluit, 1979; Monger et al., 1991). These events may correlate in part with post-Sonoma (Orogeny) magmatic arc and successor-back-arc basin development in the southern USA Cordillera (Ingersoll, Chapter 11, this volume). However, the sedimentary record indicates that it was not until the late Early Jurassic to early Middle Jurassic that fundamental changes in the Cordilleran landscape occurred, with collision between Quesnellia and Stikinia, and subsequent delamination of the intervening Cache Creek oceanic crust during southwest-directed obduction over Stikinia (Monger et al., 1972, 1982; Eisbacher, 1981), possibly as early as Aalenian (Ricketts et al., 1992). Blueschist metamorphism of Cache Creek rocks as young as Early Jurassic is dated at 173.7 Ma and was followed

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Table 2 Principal characteristics of allochthonous and pericratonic terranes in the Canadian Cordillera. The modern ocean plates are included for completeness. Abbreviations are those used in Figure 1b. The descriptions are mainly from Monger et al. (1982), Monger (1989), Jones et al. (1986), Gabrielse et al. (1991), and other references cited herein. Allochthonous & Pericratonic Terranes

De¢ning characteristics

Alexander, AX

Possible composite of Precambrian (?) Paleozoic and Mesozoic volcanic, clastic sedimentary and limestone. Associated with WR in the Paleozoic. Translated from lower latitudes Mississippian to Lower Jurassic oceanic rocks, structurally complex, metamorphosed to blueschist facies. Overlain by Jurassic to L. Cretaceous sediments (Cayoosh Assemblage) Carboniferous to Jurassic rocks having oceanic afﬁnities, including radiolarian chert, ophiolite; blueschist metamorphism Pericratonic; Precambrian to Devonian platform carbonates having afﬁnities with the Western Canada miogeocline Triassic-Jurassic arc rocks; Triassic to Mid-Cretaceous carbonates & clastics Composite terrane; Upper Mesozoic ﬂysch and melange, Paleogene ﬂysch and volcanic rocks Eocene ophiolitic layered gabbro, submarine and subaerial lava ﬂows. Located above the modern JDF Modern ocean plate. Subduction beneath North America Pericratonic; Metamorphosed Paleozoic carbonate and clastic sedimentary and volcanic rocks having afﬁnities with the western North American miogeocline Triassic oceanic rocks, including ophiolites, overlain by Jurassic to Cretaceous clastic assemblages Pericratonic. Lower Proterozoic, high-grade metamorphic rocks Modern ocean plate. Separated from North America by Queen Charlotte transform Jurassic-Cretaceous continental margin chert, sandstone, conglomerate, volcanics, and metasediments; melange. Located above the modern JDF Devonian to Upper Paleozoic volcanics and carbonates; Triassic–Jurassic volcanic, plutonic and associated sedimentary rocks. Faunas indicate terrane originally at lower latitudes Upper Paleozoic mudrock and coarse-grained clastics, overlain by Alpine-type ultramaﬁc rocks and carbonates. Permian fusulinids indicate paleolatitudes equivalent to northern Mexico or southern California The largest terrane in the Canadian Cordillera. Mainly Devonian to Jurassic volcanic and plutonic rocks and associated sediments. Faunas indicate derivation from lower latitudes Structurally complex, Upper Paleozoic to Lower Mesozoic volcaniclastics and basalt, limestone, and ﬂysch Composite terrane, Upper Paleozoic to Jurassic magmatic complexes, limestone, and clastic sediments. Associated with AX in the Paleozoic. Translated from lower latitudes Transitional continental crust of metamorphosed Cretaceous to Paleogene sediments. Translated northwards along the Queen Charlotte-Fairweather transform Pericratonic, includes part of Omineca Belt; variably metamorphosed Paleozoic and possibly Precambrian to Jurassic sedimentary, volcanic and granitic rocks; similarities to K

Bridge River, BR Cache Creek, CC Cassiar, CA Cadwallader, CD Chugach, CH Crescent, CR Juan De Fuca, JDF Kootenay, K Methow, M Monashee, MO Paciﬁc Ocean, PO Paciﬁc Rim, PR Quesnellia, Q Slide Mountain, SM

Stikinia, ST Taku, T Wrangellia, WR Yakutat, Y Yukon-Tanana, YT

by rapid exhumation of the metamorphosed subduction assemblage (Mihalynuk et al., 2004). In their interpretation of deep reﬂection-seismic proﬁles across the northern Cordillera, Evenchick et al. (2005) show Stikinia overlying a westward-tapering wedge of North American crust. The upper contact of the crustal wedge is inferred to be the accretion surface resulting from the interaction between Stikinia, Cache Creek, and ancestral North America that includes pericratonic terranes (Figure 5). In north-central British Columbia, ﬂexure of Stikinian crust beneath the Cache Creek terrane load, in concert with thermoisostatic cooling of Stikinia, resulted in a foredeep — Bowser Basin. Age constraints for this event are cemented where King Salmon fault and associated crustal-scale structures are cut by 160–172 Ma postkinematic plutons (Woodsworth et al., 1991). Collision of the contiguous Intermontane terrane with the North American margin took place about the same time (Monger et al., 1982; Murphy et al., 1995). On the basis of U/Pb dates from Kootenay Arc and Caribboo Mountains, Murphy et al. (1995) suggested overthrusting of the autochthon by Quesnellia in Mid-Toarcian time, and southwest-vergent deformation until the end of the Aalenian, the latter age bracket corresponding with Cache Creek obduction over Stikinia (Ricketts et al., 1992). Unterschutz et al. (2001) on the other hand contended that Quesnellia in southern British Columbia had depositional links with North America during the Late Triassic, based on Nd isotopic and geochemical data

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Figure 3 Chart showing the space--time relationships among major sedimentary basins of the Canadian Cordillera, the mor photectonic belts, associated terranes--superterranes and their accretion to North America (adapted from Monger, 1989; Yorath, 1991; and other references cited herein). Large arrows indicate the generalized polarity of sediment £ux into the basins.

(REEs) that indicate both a primitive sediment source (volcanic arc) and a more evolved source from the craton. This also implies that Quesnellia may be pericratonic and/or that accretion to North America was strongly diachronous. Additional sedimentological and Nd/U-Pb isotope data from the Nicola horst in southern British Columbia, indicate that mid-crustal Precambrian rocks there have continental afﬁnities, supporting a pericratonic origin for part of Quesnellia (Erdmer et al., 2002). The Quesnel–Cache Creek–Stikinia interaction is linked thermally and mechanically to metamorphism and tectonic thickening of the Omineca belt (the ‘‘core zone’’ of the eastern Foreland belt; Brown et al., 1986). In the foreland basin these events correlate with deposition of the Kootenay-Fernie clast wedge (Figure 3) (Stockmal et al., 1992; Miall et al., Chapter 9, this volume). Thickening, in part is also the result of tectonic onlap of Slide Mountain and Quesnellia over the western Omineca belt during contraction (Gabrielse et al., 1991).

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Figure 4 A summary of the main provenance characteristics for each sedimentary basin (see text of relevant basins for sources of information). Black arrows indicate the generalized polarity of sediment supply into the basins.Terrane abbreviations are from Table 2. Additional abbreviations are: CB, Coast belt; Casc. Arc, Cascade arc;Volc., volcanic source; N. Calif., northern California; OM, Omineca belt.

In north-central British Columbia the stratigraphic record of these events lies within the Bowser Basin overlap assemblage. Farther south, the Tyaughton–Methow trough is preserved at the intersection of the Quesnel–Cache Creek–Stikine terrane amalgam, with two smaller terranes, the Bridge River and Cadwallader terranes which have oceanic and magmatic arc afﬁnities, respectively. Middle Jurassic strata in Tyaughton–Methow basin directly overlie Bridge River and Cadwallader rocks and constitute an overlap assemblage (Kleinspehn, 1985; Umhoefer et al., 2002). McClelland et al. (1992) suggested that Bowser and Tyaughton–Methow basins were linked both depositionally and kinematically by an oblique dextral-slip fault system during the Late Jurassic to Early Cretaceous.

3.2. Insular superterrane In Canada, the Alexander terrane and Wrangellia are the main crustal elements comprising the amalgam of the Insular composite or superterrane (Table 2; Figures 2 and 3): Alexander terrane consists of Paleozoic (and ?Precambrian) arc, and possibly rift-related Mesozoic assemblages; Wrangellia itself is a composite terrane containing a variety of arc-related assemblages that extend from southern Alaska to beneath Vancouver Island (Jones et al., 1986). The Peninsula terrane in southern Alaska is also amalgamated with Wrangellia (Jones et al., 1986). Dated granite intrusions that stitch basement rocks from Alexander and Wrangellia terranes indicate that the two have probably been together since at least the Late Pennsylvanian (Monger, 1989). The earliest signals of accretion of the Alexander–Wrangellia–Peninsula terrane amalgam to the North American margin include 175 Ma metavolcanics in southeast Alaska, that overlap the Alexander and YukonTanana terranes (Gehrels, 2001). The older age limit for amalgamation of Insular and Intermontane superterranes is recorded by the Upper Jurassic–Lower Cretaceous Gravina–Nutzotin overlap assemblage in the northern Cordillera, and the Gambier overlap assemblage, that includes the Dezadeash and Gambier basins in central and southern British Columbia (Monger et al., 1982, 1994). These basins (Figure 2) formed on the inboard margin of the Insular Superterrane (McClelland et al., 1992); whether they were a continuous or discontinuous retroarc or forearc basin is still debated (McClelland et al., 1992). Nevertheless, there seems little doubt that they were

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Figure 5 Sketch from part of the SNORCLE re£ection seismic pro¢le (line 2b) showing the primary terrane components and main faults from the northwestern edge of Bowser basin north to the Cache Creek--Quesnellia terrane boundary (from Evenchick et al., 2005). Details of the seismic pro¢le are provided by Snyder et al. (2002) and Cook et al. (2004). The accentuated part of line 2b in the accompanying simpli¢ed geology-terrane map corresponds to the sketched portion of the seismic pro¢le shown here; vibration point 5,000 is located at the south end of the sketch.

associated with accretion of the Insular Superterrane to North America. In contrast, Wrangell Mountains and Queen Charlotte forearc basins formed on the outboard margin of the Insular Superterrane (Figure 2). Monger et al. (1982) and others maintain an Early Cretaceous time for the ﬁnal docking of Intermontane and Insular superterranes, which were subsequently stitched together by Cretaceous–Paleogene granite plutons in northwestern and southwestern British Columbia. The accretion boundary along the eastern margin of Insular Superterrane (Alexander–Wrangellia terranes) is marked by a Mid-Cretaceous, west-verging thrust belt that extends from southern Alaska to the northwestern USA (Monger et al., 1982; McClelland et al., 1992; McClelland and Mattinson, 2000). Emplacement of the thrust belt also marked the collapse of the inboard Gravina–Nutzotin–Dezadeash–Gambier collisional, arc-related basins. Both the basinal strata and Alexander– Wrangellia basement rocks form the footwall to the east-dipping thrust system. In southeastern Alaska and northwestern British Columbia the thrust system hanging wall contains elements of Taku terrane, and a metamorphic assemblage that is correlated with Yukon-Tanana terrane rocks based on Nd-Sr isotope and detrital zircon characteristics (McClelland et al., 2000; Saleeby, 2000; Gehrels, 2001). The intervening welt of plutonic and metamorphic rocks, the Coast belt, contains mostly Cretaceous-Tertiary granitic rocks that evolved initially as a magmatic arc above an east-dipping subduction zone. Exposed granulite-grade metamorphic rocks in the Coast belt indicate signiﬁcant crustal thickening and subsequent uplift of as much as 25 km. Mid-Cretaceous crustal shortening along the eastern margin of the Insular Superterrane was approximately coeval with accumulation of the Blairmore clastic wedge in the Alberta foreland Basin (Figure 4) (Stockmal et al., 1992; Miall et al., Chapter 9, this volume). In California at about this time, the Great Valley forearc basin was coupled to the Franciscan subduction complex above an east-dipping subduction zone (Ingersoll, Chapter 11, this volume). Separating the west-verging thrust belt and Late Cretaceous to Eocene plutons along the western Coast belt, is the Coast shear zone that extends about 700 km into southern Alaska and British Columbia. The steeply dipping

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shear zone developed after accretion of Insular Superterrane to North America and may have accommodated some of the transcurrent terrane displacements relative to the stable craton (McClelland et al., 2000). Debate continues about the timing of accretion events and about the crustal position of the basins involved prior to their demise following superterrane accretion. For example, the inboard Gravina–Nutzotin basins have been regarded as: (i) marginal to a Late Jurassic ocean basin (Monger et al., 1982); (ii) a Late Jurassic rift basin (e.g., Van der Heyden, 1992); (iii) a Late Jurassic dextral transtensional basin that was transported northwards (e.g., McClelland et al., 1992); (iv) a back-arc basin associated with oblique sinistral convergence (Monger et al., 1994); and (v) part of the Alexander–Wrangellia system that originated near Mexico (e.g., Irving et al., 1985). Kapp and Gehrels (1998) evaluated each hypothesis in light of zircon U-Pb isotope data and concluded that the Gravina–Nutzotin basin likely formed in a rift or transtensional setting between the composite Wrangellia– Alexander terrane and the Late Jurassic margin of North America that extended from Central America to Alaska. Trop et al. (2002) observed that basin subsidence and ﬁlling of the outboard Wrangell Mountains and Queen Charlotte basins is the reciprocal of the inboard Gravina–Nutzotin–Dezadeash basins, where the period of deposition in the latter (Late Jurassic–Early Cretaceous) corresponds to a signiﬁcant hiatus in the Wrangell Mountains-Queen Charlotte basins. Likewise, Early to Late Cretaceous deposition in the basins outboard of Insular superterrane corresponds to non-deposition in the inboard Gravina–Nutzotin–Dezadeash basins. If correct, the interpretation implies that basin development on both the eastern and western margins of the Insular superterrane were linked by the dynamics of superterrane accretion. Beginning in the Early Cretaceous, regional shortening across Bowser Basin in north-central British Columbia produced the Skeena fold belt, that was broadly coeval and possibly dynamically linked to similarstyled deformation in the Western Canada foreland belt; deformation may also be linked to Coast belt development (Evenchick, 1991, 2001). Northwestward tectonic shortening in the fold belt is as much as 160 km (Evenchick, 1991). The resulting (piggyback) thrust-top basin, Sustut basin, evolved in concert with, and was subsequently deformed by the Skeena deformation. Mid-way through the Early Cretaceous, regional uplift and erosion of the Intermontane belt and OminecaForeland belts, coincided with major sub-Hauterivian/Barremian unconformities in Tyaughton–Methow and Bowser basins, and beneath the Blairmore clastic wedge in Alberta Basin (Figures 3 and 4) (Eisbacher, 1981; Yorath, 1991; McClelland et al., 1992). This event is also signaled by a signiﬁcant decrease in intrusion and volcanism farther west (Armstrong, 1988) and by a decrease in plate velocities, in particular the westward motion of the North America craton relative to Farallon plate and Kula plate (Engebretson et al., 1985). Sediment derived from uplift and plutonism in Omineca belt was for the ﬁrst time transported westward into Sustut and Skeena basins (ﬁrst appearance of white micas), and Tyaughton–Methow trough (Figure 4). MidCretaceous plutonism in the southern Coast belt also shed sediment into Tyaughton–Methow trough. Arc-related sedimentation continued in Gambier Basin. Continued crustal shortening across the orogen during the Campanian through Paleocene revived uplift and erosion of the Foreland belt providing sediment for the Belly River and Paskapoo clastic wedges (Figure 3). Signiﬁcant uplift in the Coast belt and continued tectonic shortening across the inverted Bowser Basin also resulted in the reversal of sediment transport to Sustut Basin, and westerly oriented sediment dispersal into the newly subsiding Georgia Basin that overlaps Wrangellia and the southern Coast belt. Tectonic shortening also continued across Tyaughton–Methow trough. This phase apparently was driven by accretion of three relatively small terranes: Chugach terrane in northern British Columbia and Alaska (Jones et al., 1986), and Paciﬁc Rim and Crescent terranes on southwest Vancouver Island. Both Paciﬁc Rim and Crescent terranes are presently wedged beneath Wrangellia, above the subducting Juan de Fuca plate (Figure 2) (Hyndman et al., 1990; Cook et al., 1991). Thus, in Late Cretaceous–Paleogene times, strain was distributed across the entire width of the Cordilleran orogen. During the Early Tertiary deposition continued along the west margin of Insular belt in Queen Charlotte and Georgia basins. Toﬁno Basin that encompasses most of the continental shelf west of Vancouver Island, contains Eocene to Recent marine clastic sediments that overlie, from east to west, Paciﬁc Rim terrane, Crescent terrane, and the modern accretionary prism. Inland, non-marine sedimentation inﬁlled fault-bounded basins and troughs that commonly were associated with regional, north-northwest-striking strike-slip faults (e.g., Northern and Southern Rocky Mountain trenches, Tintina fault and Denali fault). These fault systems also dismembered older basins, such as Dezadeash Basin which is truncated and offset 300–400 km by the Denali fault system (Eisbacher, 1976; Lowey, 1998).

3.3. Magmatism, deformation, and relative plate motions Ground-breaking studies of Coast belt magmatism by Armstrong (1988) revealed periods of intense intrusive and volcanic activity punctuated by intervals of relative magmatic quiescence. Furthermore, it has been demonstrated

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that these magmatic ‘‘cycles’’ exhibit a degree of correlation with periods of high or increased velocities of the Farallon, Kula, and Paciﬁc plates relative to the North American craton (Engebretson et al., 1985; Woodsworth et al., 1991; Ward, 1995). In the Canadian Cordillera, inception and/or changes in the rate of basin subsidence, accompanied by high sediment ﬂux, also appear to be broadly sympathetic with the acme of magmatic activity, and with major periods of supracrustal deformation in the Western Canada foreland fold and thrust belt, Skeena fold belt, and the Coast Mountains thrust belt (Figures 3 and 4).

3.4. The modern plate boundary The present relative motion of North America is toward the Paciﬁc Ocean, relative to mantle hotspots (Riddihough and Hyndman, 1991); in the last 180 Ma the craton has moved through about 701 of longitude (Engebretson et al., 1985). Between northern California and Alaska the modern North American plate margin is bounded by two major transform faults, two subduction zones and two oceanic plates (Figure 6). From the south, the dextral San Andreas Fault separates the North American from the Paciﬁc plate. The Mendocino triple junction marks the transition to the Juan de Fuca plate that extends north to about the southern tip of Queen Charlotte Islands; the Juan de Fuca Plate is moving eastward and is being subducted beneath North America at up to 46 mm/year relative to North America (Cascadia subduction zone). Deep reﬂection seismic shows that the top of the downgoing Juan de Fuca plate is about 30 km beneath western Vancouver Island (Yorath et al., 1985; Hyndman et al., 1990; Cook et al., 1991). The Juan de Fuca plate is separated from the Paciﬁc plate by enechelon, northeast-trending spreading ridges (Juan de Fuca and Explorer ridges) and transform faults. A triple junction offshore Vancouver Island is hypothesized to be evolving at the junction of Juan de Fuca Ridge and Nootka fault (Rohr and Furlong, 1995). In this model Explorer microplate, born about 5 Ma, accounts for strain partitioning between Juan de Fuca plate and the Queen Charlotte transform. North of Vancouver Island, the Paciﬁc plate, and the much smaller Yukatat terrane have a strong northward component of motion, relative to North America, that is taken up by the Queen Charlotte-Fairweather dextral transform fault. North of the British Columbia–Alaska border, this motion is almost orthogonal such that Paciﬁc crust is being subducted beneath Alaska at the Aleutian Trench.

3.5. The ‘‘Baja BC’’ debate Whereas it is generally accepted that the Cordilleran terranes were laterally mobile, there is still considerable debate over how far they traveled before their accretion to North America. There are essentially two schools of thought: one that posits travel from afar, up to 5,000 km and usually from warmer southern climes; and the other that claims the terranes have traveled less than 1,000 km. The debate hinges on the interpretation of paleontological and paleomagnetic data and terrane tracks that help determine paleolatitudes, and piercing points that aim to correlate stratigraphy, sediment provenance, or magmatism across faults. The Baja BC controversy in particular has been reviewed by Cowan et al. (1997). Inferred paleolatitudes of some Late Paleozoic and Mesozoic faunal provinces in Cordilleran terranes are anomalous when compared with provinces in autochthonous strata. For example, Permian fusulinids (Monger and Ross, 1971) and low latitude Triassic corals and molluscs (Tozer, 1982) are found much farther north in strata of accreted terranes (e.g., Slide Mountain terrane) than in strata on the craton, and indicate signiﬁcant northwards displacement. In contrast, latest Jurassic and Cretaceous molluscan and radiolarian faunas from the Insular superterrane indicate middle to high paleolatitudes and therefore minimal northward displacement (Haggart and Carter, 1994). Paleomagnetism data collected over the last half century (see the review by Irving and Wynne, 1991) indicate signiﬁcant latitudinal displacements for (summarized by Keppie and Dostal, 2001): Baja Alaska, comprising Wrangellia, Peninsular, and Chugach terranes (B5,000 km); and Baja BC (including Insular superterrane B3,000 km, Stikinia B2,000 km, and Quesnellia B1,000 km). Insular and Intermontane superterranes appear to have amalgamated by Mid-Cretaceous time and according to the paleomagnetism data this amalgam was located about 2,000 km south of its present position (Irving and Wynne, 1991). However, even here there are signiﬁcant discrepancies, as is the case for the southern block of the Methow terrane (displaced northwards 1,800 +/�500 km) compared with the northern block of the Methow terrane that was displaced B3,000 +/�450 km; syntectonic remagnetization of the southern block has been posited as an explanation for the contradictory data sets (Enkin et al., 2002). Structural tilting of crustal blocks has also been shown to generate discrepancies in paleomagnetically determined terrane paleolatitudes. For example, Neogene tilting of parts of southern Alaska and British Columbia mean that coastwise displacements of less than 1,000 km are required to explain the discrepancies in paleolatitude (Butzer et al., 2004). Very little lateral displacement of Baja BC has taken place since the Mid-Tertiary (Irving and Wynne, 1991).

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Figure 6 The modern plate boundary of western Canada and northwestern USA. Adapted from Monger (1989) and Riddihough and Hyndman (1991).

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Sediment provenance methods that are used to unravel the source rock links between terranes and western North America include whole-rock petrography, zircon U-Pb ages, REE isotopes and geochemistry (e.g., Garver, 1992; Garver and Brandon, 1994; Gehrels et al., 1995; Kapp and Gehrels, 1998). Unfortunately, arguments relating to terrane displacement based on these data frequently conﬂict not only with paleomagnetism data, but between the provenance data sets themselves. For example zircon ages from Mesozoic strata in Methow, Gravina, Queen Charlotte, and Georgia basins generally indicate derivation from source rocks that require little separation from their present positions (Kapp and Gehrels, 1998; Mahoney et al., 1999; DeGraaf-Surpless et al., 2003). This interpretation depends in part on a comparison of Precambrian zircons with particular segments of North American crystalline basement (e.g., Mahoney et al., 1999). However, these interpretations are challenged by Housen and Beck (1999) and Keppie and Dostal (2001), who suggest that components of the Precambrian zircon population could also be derived from northern Mexico and therefore support the paleomagnetic data. Keppie and Dostal (2001) also propose a test that involves correlation of plume magmatism as piercing points, for example where latest Cretaceous magmatism in northern Stikinia is correlated with the Yellowstone hotspot. Clearly, adjudication of this debate in a way that is acceptable to the broader geological community is not yet possible, but its resolution is important because there are signiﬁcant implications for paleogeographic reconstructions.

4. Sedimentary Basins Associated with Intermontane Superterrane Sedimentary basins directly related to Cordilleran terrane accretion are treated in three sections that follow approximately the accretionary events from east to west: (1) Intermontane Superterrane accretion, (2) Insular Superterrane accretion, and (3) basins originating from Cenozoic plate-tectonic processes leading to the modern plate boundary.

4.1. Whitehorse trough 4.1.1. Tectonostratigraphic foundations The sedimentary ﬁll of Whitehorse trough provides evidence of the oldest (Early Jurassic and possibly latest Triassic) linkages between Stikinia, Cache Creek terrane, Quesnellia, and North America (Monger, 1989). The basin is a structurally shortened and largely fault-bound remnant of an Early Jurassic, northeast-facing forearc basin that overlies both Stikinia and Cache Creek terranes (Eisbacher, 1985; Mihalynuk et al., 2004). It is about 500 km long and is located between the King Salmon fault in the south and west, and Nahlin fault (Figure 7). Outliers that unconformably overlie Stikinian arc rocks also are present south of King Salmon fault. Whitehorse trough formed above a southwest-dipping subduction zone in which Cache Creek oceanic crust was thrusted beneath Stikinia. Blueschist metamorphism between 191 Ma and about 177 Ma (Mihalynuk et al., 2004) during subduction was coincident with the earliest period of sedimentation in the forearc basin. Subsidence was sufﬁcient to accommodate more than 4,000 m of Sinemurian to Bajocian clastic, mostly deep-water sediments (Laberge Group) (Eisbacher, 1974c, 1985; Tempelman-Kluit, 1979). The Laberge Group overlies at least 4,000 m of Stikinian arc-related rocks (Lewes River Group). Uplift in Middle Jurassic time resulted in the demise of the forearc, and concomitant basin shallowing with deposition of paralic, deltaic, and ultimately alluvial sediments in the upper part of the Laberge Group (Tanglefoot Formation) and the disconformably overlying Tantalus Formation. Progradation of these Upper Jurassic–Lower Cretaceous successor-basin assemblages is correlated with the ﬁnal stages of Stikinia–Cache Creek terrane accretion (Dickie and Hein, 1995). Forearc basin subsidence in the Early Jurassic is delineated by the abrupt transition from lowest Triassic to Hettangian shallow-marine facies, to the relatively deep-water Laberge Group (Figure 7). The shallow water facies include bouldery coastal fan deltas that interﬁnger with sandy delta platform and shelf facies and local bioherms. The Sinemurian through Bajocian portion of the Laberge Group consists primarily of facies developed in slope, slope-apron, and submarine-fan environments (Dickie and Hein, 1995; Hart et al., 1995; Johannson et al., 1997). 4.1.2. Provenance linkages Signiﬁcant changes in clast composition through the Laberge succession reﬂect fundamental changes in tectonic setting with respect to terrane accretion, and evolving source rocks (Figure 8). Petrofacies analysis of Sinemurian to Pliensbachian strata indicate that most sediment was derived from unrooﬁng of older Stikinian arc rocks, and Early Pliensbachian volcanic rocks (Johannson et al., 1997). Uplift and incision of Pliensbachian granitic plutons (U-Pb B186 Ma) was rapid enough to supply large volumes of sediment during Pliensbachian deposition. U-Pb

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Figure 7 Map of Whitehorse trough separated from Cache Creek terrane (CC) by Nahlin fault, from Stikinia (ST) by the King Salmon fault, and in the north from Quesnellia (Q) by the Teslin fault. Note the stratigraphic contacts at the northern, and western limits of the trough. Modi¢ed from Johannson et al. (1997) and Hart et al. (1995).

Figure 8 Schematic summary of provenance trends for the Whitehorse trough. Source terranes are shown in the left column, and unroo¢ng pathways on the right column (Laberge Group). The fundamental shift in clast composition in Middle Jurassic strata re£ects the uplift and exhumation of Cache Creek terrane rocks following their earlier Jurassic blueschist-grade metamorphism. Adapted from Johannson et al. (1997) and Mihalynuk et al. (2004).

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zircon dating of granite boulders conﬁrms the Stikinian arc as the likely source (Hart et al., 1995). Signiﬁcantly, there is no evidence for Cache Creek-derived sediment in Whitehorse trough during the Early Jurassic. Early Jurassic unrooﬁng of Stikinian rocks was coincident with blueschist metamorphism of the subducted Cache Creek terrane. Metamorphism of Pliensbachian to Toarcian Cache Creek chert has been dated at about 173 Ma (Mihalynuk et al., 2004). Rapid exhumation and unrooﬁng of the metamorphic complex supplied chert and granulite clasts to Whitehorse trough before about 171 Ma (Early to Mid-Bajocian), a date that is constrained by the ages of cross-cutting, post-kinematic plutons (Figure 7) (Mihalynuk et al., 2004). These events were likely related to the closure of the Cache Creek Ocean between Stikinia and Quesnellia.

4.2. Bowser Basin 4.2.1. Tectonostratigraphic foundations Bowser Basin, the largest contiguous basin in the Canadian Cordillera, is located between Stikine and Skeena arches and extends about 400 km parallel to the orogen (Figure 9). The signal importance of Bowser Basin lies in its provenance links to obducted, uplifted and eroded Cache Creek terrane during the ﬁnal stages of collision between Quesnellia and Stikinia. The bedrock foundations of the basin consist of the Triassic–Mid-Jurassic Hazelton volcanic assemblage, in particular the Early-Middle Jurassic Hazelton Group. Hazelton Group arc volcanism extended over much of Stikinia and began to decline in the Toarcian with subsequent thermal subsidence of Stikinian basement. The combined effects of thermal subsidence and crustal-scale loading by obducted Cache Creek resulted in widespread ﬂexural subsidence (Monger, 1977; Eisbacher, 1985; Evenchick and Thorkelson, 2005). Initial subsidence is recorded in the Upper Spatsizi Formation (uppermost Hazelton Group; Figure 10) by extensive deep-water mudrocks (Quock Mbr.) and more locally condensed, organic-rich

Figure 9 Map of Bowser and Sustut basins, and the present relationships with Cache Creek terrane,Whitehorse trough, and Stikinian basement. Generalized structural trends are based on mapped fold axes and thrust traces (modi¢ed from Evenchick, 1991, 2001; Mihalynuk et al., 2004).Whitehorse trough is separated from Stikinia by the King Salmon fault (dashed line). ST, Stikinia; CC, Cache Creek terrane; Q, Quesnellia; NA, autochthonous North America; Su, Sustut basin; HB, Hotailuh batholith.

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Figure 10 General stratigraphy and schematic of Bowser Lake Group depositional model along an approximately southwestoriented pro¢le (modi¢ed from Eisbacher, 1981; Yorath, 1991; Ricketts and Evenchick, 1991, 1999; assemblages from Evenchick and Thorkelson, 2005). The Abou and Quock members at the base of the succession belong to the upper part of the Spatsizi Formation that underlies the Bowser Lake Group. CC, Overthrust Cache Creek terrane.

shale (Abou Mbr.) that overlie a sub-Aalenian erosional unconformity (Monger et al., 1991). This (Aalenian) starved-basin stage, represented by Quock-Abou sedimentation, was interpreted by Ricketts et al. (1992) to reﬂect initial ﬂexural subsidence of Stikinia. However, Gabrielse (1991) has noted that collision may have begun as early as Toarcian based on stratigraphic relationships on Stikine arch. The Toarcian–Aalenian interval also corresponded with the period of signiﬁcant convergence and tectonic shortening between Quesnellia and North American in southern British Columbia (Murphy et al., 1995). Regionally, Bowser Basin was a west-facing foredeep that continued to subside until the Early Cretaceous. The basin accommodated a broadly regressive, southwest-prograding marine and non-marine succession, up to 3,500–4,000 m thick, that can be divided into three main components (Figure 10): (1) the starved-basin stage (Upper Spatsizi Formation) dominated by mudrock; (2) the Bowser Lake Group that makes up most of the basin ﬁll, and: (3) the Devils Claw Formation (conglomerate) that heralds deformation associated with shortening across the Skeena fold belt. The northern and eastern margins of the basin are overlain unconformably by the Cretaceous-Maastrichtian Sustut Group (Sustut Basin), a foredeep that piggybacked on the deformed Bowser Basin during Skeena fold belt contraction. The stratigraphic nomenclature of the basin has evolved such that formations are deﬁned in some areas (Tipper and Richards, 1976; Bustin and Moffat, 1983), and facies or facies assemblages in other areas (Eisbacher, 1974a; Evenchick and Thorkelson, 2005). The problem with the more formal stratigraphic schemes is that formations and members are difﬁcult to map beyond their type sections because of the structural complexity. In this synopsis, the more recent nomenclature of Evenchick and Thorkelson (2005) is used. Across the basin, the overall trend toward southwest-directed regression and regional ofﬂap is depicted by Eisbacher (1974a, 1981) as a delta system with associated ﬂuvial, slope, and basin-ﬂoor facies (Figure 10). However, the northern basin margin for much of the Bathonian to Oxfordian-Kimmeridgian interval was also characterized by coarse-grained facies (Eaglenest and Muskaboo Creek assemblages) that include Gilbert and

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braid deltas that prograded onto narrow, tectonically active sand-dominated shelves (Ricketts and Evenchick, 1999, 2007). Basinward, slope facies (Todagin assemblage) and basin-ﬂoor fans (Ritchie-Alger assemblage) were depositionally linked to the shelf — fan-delta systems by shelf-break gullies and gravel-ﬁlled submarine channels (Figure 11). Gravel mostly bypassed the shelves during sea-level lowstands. Submarine channels incised into the slope deposits locally occur as stacked and overlapping complexes, that are ﬁlled with gravel deposited as cohesive and cohesionless debris ﬂows (Ricketts and Evenchick, 1991). Continued regression during the latest Jurassic to Early Cretaceous resulted in extensive delta plains that contained swamp, lacustrine, ﬂoodplain, and distributary and high-sinuosity channel facies (Groundhog– Gunanoot and related assemblages) (Eisbacher, 1981; Bustin and Moffat, 1983; MacLeod and Hills, 1990). The transition from the coarser-grained, predominantly marine assemblages farther north to the delta-plain assemblages may correspond to a decrease in the rate of foreland-driven subsidence and either: (i) a decrease in surface topography associated with the over-thrust Cache Creek, or (ii) the northwards retreat of exposed Cache Creek source rocks toward their present position in the hanging wall of King Salmon fault (Figure 9). The non-marine Devils Claw Formation, up to 1,000 m thick and containing 30–80% conglomerate (Evenchick and Thorkelson, 2005), is the youngest unit in the Bowser Lake Group. It abruptly but conformably overlies Groundhog–Gunanoot strata. Devils Claw strata herald a return to high-energy alluvial fan — bedload stream conditions (Eisbacher, 1981) that are equated with initial Skeena fold belt deformation in the late Early Cretaceous, and the cannibalizing of older Bowser Basin strata. Devils Claw strata in turn are unconformably

Figure 11 Inferred depositional setting for Callovian to Oxfordian conglomerate-dominated assemblages on the northern Bow ser basin margin. During this interval shelves were probably narrow and tectonically active, commonly overlain by large Gilbert deltas. Sediment £ux was high because of proximity to the obducted Cache Creek block. Coarse-grained sediment frequently bypassed the shelf during both highstands and lowstands of sea level, eventually being deposited as slope apron and basin £oor submarine fans. Adapted from Ricketts and Evenchick (1991, 1994).

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overlain by units of Sustut Basin. South toward Skeena Arch, approximately correlative Hauterivian to Aptian strata in the Skeena Group, of paralic and neritic origin (Tipper and Richards, 1976), overlie a transgressive surface on the Hazelton and Bowser Lake groups. These in turn are overlain by upper Lower Cretaceous subaerial volcanics and marine to non-marine sediments. Early studies of coaliﬁcation in the northwestern part of Bowser Basin indicated anthracite to meta-anthracite levels that Bustin and Moffat (1989) equated to about 5,000 m of burial and an average heat ﬂow of 80 mW/m2. However, recent evaluation of coal rank and oil preservation potential in the northern two-thirds of the basin indicate that the regional coal ranks lie within the oil–gas generation–preservation window (Osadetz et al., 2003). These recent discoveries have sparked renewed interest in Bowser Basin as a possible petroleum province. 4.2.2. Provenance linkages By far the most abundant clast-type making up Bowser Basin rocks is radiolarian chert in sandstone and conglomerate; W90% in many conglomerate outcrops. Conglomerate chert clasts are spherical to oblate and very well rounded, indicating a substantial degree of mechanical abrasion during transport; recycling of Bowser Basin sediment was also likely (Ricketts and Evenchick, 2007). Limestone clasts (from the Cache Creek terrane) are less common (o5%); metamorphic detritus is rare. The only conceivable source for such large volumes of chert is the oceanic Cache Creek terrane (Souther and Armstrong, 1966; Eisbacher, 1981; Gabrielse, 1991; Evenchick and Thorkelson, 2005). Recognition of Alpine-type spinals in the heavy mineral sediment fraction supports this interpretation (Cookenboo et al., 1997). Eisbacher (1981) also noted that the underlying Hazelton rocks provided relatively little sediment to the basin. The chert petrofacies provides a critical link among the lithosphere-scale processes of Cache Creek–Stikinia collision, overthrusting, and subsidence of the Bowser Basin. Notably, there is no evidence that the Omineca belt supplied sediment to Bowser Basin; this situation changed during formation of the overlying Sustut Basin (Figure 4).

4.3. Sustut ‘‘piggyback’’ Basin 4.3.1. Tectonostratigraphic foundations Whereas the Devils Claw Formation (uppermost Bowser Basin) heralds initial deformation in the Skeena fold belt, the overlying Sustut Group records uplift, erosion and cannibalizing of the older Bowser Basin deposits, and direct involvement of Sustut strata in Skeena fold belt contraction during the Mid- to Late Cretaceous (Figures 9 and 10). Sustut Basin is interpreted to be the foreland basin to Skeena fold belt (Evenchick, 1991). The southwest part of the basin hosts the triangle zone (Evenchick and Thorkelson, 2005), such that the basin has ‘‘piggybacked’’ with the frontal thrust. Note that Skeena fold belt contraction was driven by Insular superterrane accretion outboard (west) of the Stikinia–Cache Creek — North American amalgam, in contrast to Bowser Basin which was associated with Stikinia–Cache Creek terrane accretion. Upwards of 2,000 m of strata are divided into the Tango Creek (Barremian or Lower Albian to Upper Campanian) and Brothers Peak formations (Upper Campanian to Lower Maastrichtian) (Evenchick, 1991). Each formation represents a different stage of basin subsidence attending Skeena deformation (Eisbacher, 1974b, 1981). Basal Tango Creek conglomerate constitutes a pediment that unconformably overlies deformed Bowser Basin strata and Stikinian Hazelton volcanics (Eisbacher, 1974c). Fluvial and lacustrine deposits low in the formation accumulated in southwest-ﬂowing drainage basins, changing to northeast ﬂowing in the topmost strata (Figure 10). The coarse-grained Brothers Peak Formation unconformably overlies the Tango Creek, and heralds further changes in basin conﬁguration. Alluvial fan and ﬂuvial deposits accumulated where ﬂow mostly was directed southeast along the basin axis. The basin axis at this time appears to have been strongly controlled by the predominantly southeast-trending Skeena folds and thrust faults. Tango Creek and Brothers Peak deposition overlap in time with the Blairmore and Paskapoo-Belly River clastic wedges, respectively; Evenchick (1991, 2001) has suggested that the deformation in Skeena fold belt and the Western Canada foreland belt were dynamically linked. 4.3.2. Provenance linkages There is a fundamental provenance shift in deposits of Sustut Basin compared with Bowser Basin that corresponds to different basin dynamics and regional tectonics. Well-rounded chert clasts derived from uplifted Bowser Basin predominate in the Tango Creek and Brothers Peak formations. Eisbacher (1974b) also suggests that some Brothers Peak sediment was cannibalized from Tango Creek rocks during Skeena fold belt deformation.

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Unrooﬁng of the Omineca belt rocks provided sediment types not seen before in the northern Intermontane Superterrane, although they are encountered in Albian and younger Jackass and Pasayten groups in Tyaughton–Methow basin (Garver, 1992). Quartz clasts, locally up to 50%, abundant muscovite and foliated granite clasts were derived from high-grade metamorphic complexes and plutons, respectively. Andesite clasts in the Brothers Peak Formation indicate that uplift and erosion at this time also involved Stikinian bedrock (Eisbacher, 1981).

4.4. Tyaughton–Methow basin 4.4.1. Tectonostratigraphic foundations Remnants of the Tyaughton–Methow basin composite (Tyaughton Basin, Methow Basin) in southwest British Columbia are sandwiched between Intermontane superterrane and the eastern margin of the Coast belt (Figures 2 and 12). The basin is signiﬁcant because: it helps bracket the timing of Intermontane and Insular superterrane interaction; it deﬁnes the timing of accretion of Bridge River–Cadwallader–Methow terranes; and it provides one of the earliest source rock links to unroofed plutonic rocks in the eastern Coast belt (Kleinspehn, 1985; Woodsworth and Monger, 1991; Monger et al., 1994; Riesterer et al., 2001; Umhoefer et al., 2002; DeGraaf-Surpless et al., 2003). It is still uncertain whether the faulted segments of the basin were once part of a contiguous basin or were separate depositional entities; they are presently separated by about 100 km of dextral displacement along the Fraser– Straight Creek fault. In the North Cascades (USA), Late Jurassic volcanic arc rocks (Newby Group) are linked to the western margin of Intermontane superterrane, and underlie Methow Basin strata (Mahoney et al., 2002). The arc rocks are intruded by 152.8 Ma granodiorite, indicating that Methow terrane was linked to Intermontane superterrane at least by the Late Jurassic. A minimum age for linkage between the Insular superterrane and Bridge River– Cadwallader–Methow terranes is provided where these terranes are thrust over Coast belt rocks and where the thrust imbricates are intruded by B93 Ma plutons (Journeay and Friedman, 1993). Correlation of the Harrison Lake Formation and Cayoosh Assemblage in southeastern British Columbia further suggest a Middle Jurassic overlap assemblage linking Wrangellia (part of the Insular superterrane amalgam) and Methow terranes (Journeay and Mahoney, 1994). Tyaughton–Methow strata form an overlap assemblage on Mesozoic oceanic and arc-related rocks composing Bridge River, Cadwallader, and Methow terranes. The three terranes were amalgamated by the Callovian, based on correlation of the Relay Mountain Group (Callovian to Upper Hauterivian) and related stratigraphic units (Rusmore et al., 1988; Umhoefer et al., 2002) (Figure 12). Rusmore et al. (1988) suggested Middle Jurassic amalgamation between Bridge River and Cadwallader terranes and Intermontane superterrane based on the timing of deformation; timing that is similar to Cache Creek obduction over Stikinia. Petrofacies analysis

Figure 12 Map and generalized stratigraphic composition of segments of Tyaughton and Methow basins, associated with Bridge River, Cadwallader, and Methow terranes. Adapted from Kleinspehn (1985), Garver (1992), Umhoefer et al. (2002), DeGraaf-Surpless et al. (2003). M, Methow terrane; CAD, Cadwallader terrane; BR, Bridge River terrane; WR,Wrangellia; J-K, plutonic rocks; TCG,Taylor Creek Group; RMG, Relay Mountain Group; JMG, Jackass Mountain Group; Intermontane terranes include Stikinia, Quesnellia, and Cache Creek terrane.

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indicates that early Relay Mountain Group deposition took place in a relatively deep, forearc-like setting, west of the Tyaughton–Methow basin, from Callovian to Oxfordian time (Garver, 1992), possibly even part of the Klamath Mountains forearc (Umhoefer et al., 2002; Ingersoll, Chapter 11, this volume). The basin became shallower, and even non-marine toward the end of the Valanginian. West-derived sediment is ﬁrst recorded in the Valanginian–Hauterivian (B135–130 Ma) indicating a signiﬁcant change in paleogeography at this time with some sediment possibly sourced from the Gambier arc assemblage (Woodsworth and Monger, 1991; Umhoefer et al., 2002). Umhoefer et al. (2002) further suggested that Methow Basin at this time may have evolved to a back-arc basin located east of the Ottarasko volcanic arc. Depositional links (but not necessarily continuous depocentres) along the western margin of Intermontane Superterrane have been postulated by Eisbacher (1985) to include Tyaughton–Methow, Bowser, and Gravina–Nutzotin basins. Local unconformities separate the older overlap assemblage from the Jackass Mountain, Taylor Creek groups and related units (Hauterivian to Albian). Polymict conglomerate and volcanic-lithic wackes were deposited in a variety of submarine-fan environments (Kleinspehn, 1985) within a forearc basinal setting (Woodsworth and Monger, 1991). Uplift during Late Albian to Santonian time resulted in deposition of coarse clastic strata, including red beds (Pasayten Group and related units), that interﬁngered with volcanic ﬂows and were deposited in marginal marine, alluvial, and ﬂuvial environments across much of Tyaughton–Methow basin (Garver, 1992).

4.4.2. Provenance linkages The Relay Mountain Group sediment was derived primarily from the underlying volcanic terranes (volcanic lithic). Umhoefer et al. (2002) recognized an unrooﬁng sequence with a progressive increase in plutonic quartzfeldspar into the Oxfordian. During the Hauterivian to Albian interval, coarse volcanic and plutonic detritus was shed westwards from Intermontane Superterrane (Kleinspehn, 1985). Of particular signiﬁcance is the early provenance linkage between Insular and Intermontane superterranes during the Albian, in concert with uplift and unrooﬁng of the southern Omineca belt (Garver, 1992). Three petrofacies identiﬁed in Albian to Cenomanian strata (Upper Jackass to Lower Pasayten groups) include (Garver, 1992): a west-derived volcanic petrofacies found only in Tyaughton Basin (Taylor Creek Group); a chert petrofacies that contains minor blueschist and serpentinite detritus derived from Bridge River–Cadwallader source rocks; and an arkose petrofacies with quartz, feldspar, and muscovite derived from the Omineca belt. The arkose facies predominates in Methow Basin (strata 3–8 km thick) with only a thin unit extending to Tyaughton Basin. Detrital zircons from Mid-Albian to Santonian rocks overlying the Methow terrane also correspond best with southern Cordilleran source rocks, indicating little translation of Methow terrane relative to the North American margin (DeGraafSurpless et al., 2003). Contradistinct evidence based on paleomagnetism for Methow terrane and other components of Insular Superterrane posits upwards of 3,000 km of northwards displacement relative to the North American craton (Irving and Wynne, 1991; Keppie and Dostal, 2001; Enkin et al., 2002).

5. Basins Located along the Inboard Margin of Insular Superterrane 5.1. Nutzotin–Dezadeash–Gravina–Gambier basins 5.1.1. Tectonostratigraphic foundations Unifying elements in the debate concerning the timing of accretion and horizontal translation of Insular superterrane relative to North America, are the sediment depocentres sandwiched between Insular and Intermontane superterranes. Overlap assemblages that compose the Nutzotin, Dezadeash, Gravina, and Gambier successions (Figure 13), their provenance, and structural associations provide evidence for basin formation and the early stages of terrane collision, and eventual basin destruction following the ﬁnal stages of Insular superterrane accretion to, and tectonic overlap of the Cordilleran margin. There is no deﬁnitive evidence that there was a contiguous locus of deposition along the 2,000+ km length of the Cordilleran margin. However, there is reasonable consensus that the basins were linked in terms of Insular superterrane accretion, based on their terrane overlap geometries, timing and depositional character, and their relationships with an active, western magmatic arc (Figure 14) (Eisbacher, 1985; Monger et al., 1994; McClelland et al., 2000; Trop et al., 2002). Berg et al. (1972) incorporated the Nutzotin, Dezadeash, and Gravina successions into what is commonly referred to as the Gravina–Nutzotin belt. Connection of these inboard basins with the western margin of Bowser Basin has also been hypothesized (Eisbacher, 1985; Yorath, 1991). Links with Tyaughton–Methow basin are more tenuous although sediment derived from Gambier arc may have been transported into that basin (Garver, 1992).

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Figure 13 Generalized stratigraphic columns for Nutzotin, Dezadeash, Gravina, and Gambier basins, located on the inboard margin of Insular superterrane. Adapted from Eisbacher (1976), Rubin and Saleeby (1991), Souther (1991), Lynch (1992, 1995), Kapp and Gehrels (1998), Gehrels (2001), McClelland et al. (2000),Trop et al. (2002). Contrasting, and to some extent reciprocal events in the outboard basins are shown in Figure 16.

Figure 14 Two versions of a reconstructed Jurassic North American margin showing the relationship of the inboard Nutzotin- Gambier forearc basins to Insular and Intermontane superterranes: (A) The inboard basins together with Bowser and Tyaughton-Methow (T-M) basins are linked tectonically by a regional, oblique dextral strike--slip shear couple (from McClelland et al.,1992). Present latitudes are shown asYukon--BC and Canada--USA borders. Paleolatitudes (501N, 401N and 301N in red) also are shown on the right. AX-WR-P, Alexander--Wrangellia--Peninsula terranes (Insular superterrane). Solid black cones represent the magmatic arc. Kahiltna basin is located in north Alaska. Other abbreviations as inTable 2. (B) Eisbacher’s (1985) interpretation shows a similar disposi tion of inboard basins from Middle Jurassic to Early Cretaceous, with a more explicit link to Bowser and Tyaughton--Methow basins. Solid black arrows indicate directions of regional sediment transport. Paleolatitudes are shown by successive displacement of the Canada--USA border that depict a total of 1,500--2,000 km of displacement relative to North America over this period.

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Approximately time-equivalent terrane and island-arc suturing events that herald the Nevadan orogeny in northern California are illustrated in Figure 12 (Ingersoll, Chapter 11, this volume). Characteristics that the four basins have in common include a predominance of argillaceous- and sanddominated siliciclastic and volcaniclastic lithologies, a high proportion of sediment gravity-ﬂow deposits, and a close association with active calcalkaline and tholeiitic arc volcanism bordering the western basin margins (Figure 13). Volcanic ﬂows, agglomerates, tuffs, and welded tuffs, which were the products of strato-volcanoes and explosive volcanism (Lynch, 1992), commonly interﬁnger with the sedimentary strata. Structural imbrication and dismembering of the basin successions accompanied Mid-Cretaceous crustal shortening during the ﬁnal stages of Insular-Intermontane superterrane accretion (Journeay and Friedman, 1993). For example, the eastern boundary of Gravina Basin is overthrust by Stikine and metamorphosed Yukon-Tanana and Taku terrane rocks (Crawford et al., 1987; Rubin and Saleeby, 1991; McClelland et al., 1992).

5.1.2. Nutzotin Basin The bulk of Nutzotin Basin resides in Alaska with a small portion sneaking into the northern Canadian Cordillera. Westwards, the basin onlaps Wrangellia, whereas its eastern boundary is truncated by the dextral-slip Denali fault. More than 3,000 m of Oxfordian to Barremian, west-derived shale, volcaniclastic turbidites, and gravelly debris ﬂows compose a large-scale coarsening-upwards submarine-fan succession (Nutzotin Mountains sequence) (Berg et al., 1972; Trop et al., 2002). Minor conglomerate near the base of the succession contains Triassic limestone clasts derived from Wrangellian bedrock. Volcanic detritus was derived from the active Chisana arc bordering the west margin (Figure 15). Unrooﬁng of older and structurally deeper segments of the Chitina arc

Figure 15 Inferred paleogeographic evolution of Nutzotin (inboard; NB) and Wrangell Mountains (outboard; WMB) basins (southeast Alaska), the evolving magmatic arcs, and regional deformation associated with Insular Superterrane accretion (modi ¢ed from Trop et al., 2002). (A) The active Chitina arc lies west of Wrangell Mountains basin during the Late Jurassic--Early Cretaceous. Inception of Nutzotin basin during the Oxfordian marks the close association between the incoming Insular and Intermontane superterrane margins. (B) Arc activity jumps northeastwards during the Hauterivian with the Chisana arc forming the west margin of Nutzotin basin. (C) Thrusting, beginning in the Barremian, continues into the Mid-Cretaceous, terminating deposition in Nutzotin basin. NA, North America.

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farther west also provided igneous clasts to Lower Cretaceous conglomerates. Eastward migration of Chisana arc (relative to North America) is recorded by the gradual transition from the volcaniclastic succession to more than 3,000 m of effusive volcanics, and concomitant migration of the Nutzotin foredeep axis.

5.1.3. Dezadeash Basin The Dezadeash Formation is located northeast of Denali fault, about 300 km south of, and along strike from the Nutzotin Basin. Unlike the other three inboard basins, the preserved Dezadeash succession contains only sedimentary strata; about 3,000 m of predominantly sandy and argillaceous turbidites and other sediment gravityﬂow deposits (Eisbacher, 1976). The base and top of the succession are structurally truncated. Sediment transport was mostly toward the east and northeast from a volcanogenic source, with deposition in relatively deep-water submarine-fan lobes and channels (Lowey, 1992). The Oxfordian to Valanginian Dezadeash succession has been correlated with Nutzotin Basin rocks, wherein the Dezadeash strata are possibly the distal equivalents of a volcanogenic submarine-fan depositional system (Eisbacher, 1976). Limestone mega-boulders, derived from Triassic Wrangellian rocks, have been traced to both Nutzotin and Dezadeash successions straddling Denali fault and indicate about 370 km of strike-parallel offset (Lowey, 1998). The Kluane Schist (Kluane metamorphic assemblage of quartz-mica schist, gneiss, and serpentinite) structurally overlies, and has been posited as the metamorphosed equivalent of, the Dezadeash succession (Eisbacher, 1976). However, rare-earth element and isotopic analyses of rocks from both assemblages indicate that the Kluane Schist–Dezadeash linkage is less clear than hitherto thought (Mezger et al., 2001). Structural imbrication of the two assemblages is linked to the Mid-Cretaceous crustal shortening seen elsewhere in the Coast belt orogen (McClelland et al., 1992).

5.1.4. Gravina Basin Oxfordian to Albian strata making up Gravina Basin (southeastern Alaska) contain a complex, interﬁngering assemblage of volcaniclastic wackes, argillites, conglomerates, tuffs, and volcanic ﬂows and breccias (Berg et al., 1972; Rubin and Saleeby, 1991; Kapp and Gehrels, 1998; Gehrels, 2001). The preserved western segments of the basin unconformably overlie Wrangellia and Alexander terrane rocks, whereas eastern exposures unconformably overlie the Taku terrane (Figure 1a). Strata are separated from Stikinia and Yukon-Tanana terranes by MidCretaceous, west-verging thrusts. Metamorphism to greenschist and lower amphibolite facies pervades most of the Gravina Basin rocks (Rubin and Saleeby, 1991). The lower part of the succession contains mostly volcanic and pillowed ﬂows, pyroclastic ﬂows and tuffs that possess island-arc geochemical signatures (Rubin and Saleeby, 1991). Like Nutzotin and Dezadeash basins, basal strata also contain conglomerate with metavolcanic and limestone clasts derived from Alexander terrane, and deposited in submarine-fan channels and fan lobes. Siliciclastic turbidites are present higher in the succession. Submarine-fan development was linked to the active volcanic arc along the western basin margin. Middle Jurassic through Albian diorite and granitoid intrusions (Figure 13) apparently were shallow enough to have contributed detritus to the basin.

5.1.5. Gambier Basin Gambier arc and its associated basin are generally considered to be the southern extension of the Gravina– Nutzotin arc. Gambier Basin was located along the eastern margin of Wrangellia during the Early Cretaceous (Lynch, 1992), and subsequently was dismembered by west-vergent thrusts (Journeay and Friedman, 1993; Lynch, 1995). Geochemical trends normal to the axis of the arc suggest that Gambier magmatism developed above a west- to southwest-dipping subduction zone (Lynch, 1995). Metamorphic grade ranges from prehnite pumpellyite to amphibolite facies, and penetrative deformation is relatively intense. The Gambier Group (Roddick, 1965) is similar to the other inboard-basin successions, and contains a mix of Lower Cretaceous arc-related volcanic rocks, and volcaniclastic and siliciclastic sediments (Figure 13). The succession is in thrust and nonconformable contact with Jurassic granodiorites. Basal conglomerates (Peninsula Formation) contain boulders of granodiorite, andesite, and basalt; trough crossbedding and thin coal layers indicate local non-marine to estuarine-type environments, that pass higher in the sequence to marine sandstoneshale dominated lithologies (Lynch, 1992, 1995).

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5.1.6. Provenance linkages Sedimentary strata of the inboard basins are composed predominantly of epiclastic volcanic detritus derived from adjacent active magmatic arcs. Additional components include older sedimentary, metamorphic, and igneous rocks derived primarily from Alexander and Wrangellia terranes (Eisbacher, 1976; Rubin and Saleeby, 1991; Lynch, 1992; Monger et al., 1994; Gehrels, 2001). More precise estimates of source rock composition and age are provided by analysis of detrital zircons. For example, zircon age populations in the Gravina succession include those derived from the adjacent magmatic arc, Alexander terrane, and Paleozoic to Precambrian zircons that were likely sourced from inboard Yukon-Tanana, Stikine, or northern Californian terranes (Kapp and Gehrels, 1998). These results imply that the inboard terranes must have been close to Gravina Basin, and places some restrictions on the Late Jurassic to Early Cretaceous paleolatitude of the inboard basins (Figure 14).

6. Basins Located along the Outboard Margin of Insular Superterrane 6.1. Queen Charlotte–Wrangell Mountains basins 6.1.1. Tectonostratigraphic foundations As terranes migrate across the globe, basins associated with the leading and trailing plate edges evolve in concert with their plate-tectonic environments. Triassic and Jurassic strata in Wrangell Mountains (south Alaska) and Queen Charlotte Islands (coastal British Columbia) record the passage of the Insular Superterrane from lower latitudes northwards toward its present position. Accretion of the Superterrane to North America during the Late Jurassic–Early Cretaceous is recorded in the inboard basins (see above) and their eventual tectonic demise. These events are also recorded along the outboard margin in the Wrangell Mountains and Queen Charlotte basins. However, Trop et al. (2002) have noted that the timing of major depositional and uplift-erosional events in the outboard basins (at the basin scale) are out of phase with similar events in the inboard basins (Figures 13 and 16).

6.1.2. Wrangell Mountains Basin Uplift of Wrangellia and associated magmatic arcs in the latest Jurassic and Early Cretaceous resulted in coarse clastic sediment being shed eastward into Nutzotin–Gravina basins (Figures 2 and 15). In Wrangell Mountains Basin, this event coincides with a regional, subaerial unconformity that records the initial collision of Insular Superterrane to North America (Trop et al., 2002). Subsequent growth of Chisana arc (Hauterivian to Barremian) was accompanied by subsidence of Wrangell Mountains forearc basin and the deposition of shallowmarine clastics overlain by deeper water submarine fans. Uplift of the outboard margin during the Aptian coincided with crustal shortening in Nutzotin Basin. Renewed subsidence of the southwest-facing forearc basin and northeast migration of the magmatic arc coincided with deposition of a 3 km thick, largely upward-coarsening, submarine-fan-slope-apron succession during the Albian and Campanian.

6.1.3. Queen Charlotte Basin The Cretaceous portion of Queen Charlotte Basin (see varying deﬁnitions of the basin by Shouldice, 1971; Thompson et al., 1991; Lyatsky and Haggart, 1993) is broadly similar to Wrangell Mountains Basin (Figures 16 and 17). Lower Cretaceous strata unconformably overlie Wrangellia, where a basal shallow-marine conglomerate is the oldest indication of forearc-basin subsidence and transgression, following regional uplift that spanned most of the Late Jurassic (Lewis et al., 1991). Clast compositions reﬂect derivation primarily from local Wrangellian bedrock. Subsidence that continued until the Early Maastrichtian accommodated at least 3 km of shallow-marine and deeper water deposits. Coarse-grained, shoreface, and shallow-shelf facies in the Longarm Formation are overlain by ﬁner grained, outer-shelf, and slope facies in lower units of the Queen Charlotte Group. Continued basin subsidence and increasing water depths resulted in coarse-grained sediment bypassing the shelf and slope, and progradation of slope-apron and submarine-fan turbidites and gravelly debris ﬂows that accumulated in two distinct depocenters on Graham and Moresby islands (Honna Formation, Upper Queen Charlotte Group; Haggart, 1991; Lewis et al., 1991). Sediment transport was mostly westward. This time interval corresponds with major thrusting along the inboard margin of Insular Superterrane that resulted in the demise of the Nutzotin to Gambier belt of basins (Figure 13).

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Figure 16 Generalized stratigraphic columns for Wrangell Mountains and Queen Charlotte basins along the outboard margin of Insular superterrane. Adapted from Lewis et al. (1991) and Trop et al. (2002). Compare the overall depositional--tectonic events with those of the inboard basins in Figure 13.

6.1.4. Provenance linkages Clast compositions in both outboard basins indicate that the primary sources of sediment were: (1) the underlying Wrangellia–Alexander terranes consisting of sedimentary, metasedimentary and old magmatic arc rocks, (2) detritus derived from active arcs and local plutons, and (3) in Queen Charlotte Basin, metamorphic and peraluminous plutonic clasts that point to links with the Coast belt (Lewis et al., 1991).

7. Cenozoic Basins-Harbingers of the Modern Plate Boundary 7.1. Queen Charlotte-Georgia-Toﬁno basins Accretion of Insular Superterrane to North America was essentially over by the Mid- to Late Cretaceous. Basins that formed subsequent to these events contain Upper Cretaceous to Holocene strata that overlie Wrangellia and Alexander terranes and structurally dismembered elements of Queen Charlotte Basin (Figure 17). The basins considered here include the Eocene–Pliocene component of Queen Charlotte Basin (Figure 16), Georgia– Nanaimo basins (Figure 17), and Toﬁno Basin, all located along coastal British Columbia (Figure 2). These younger basins also onlap uplifted and dissected Coast belt plutonic and metamorphic rocks. Subsidence and deformation of the coastal basins were driven by plate-tectonic processes that evolved throughout the Tertiary. Convergent motion between Farallon plate and North America persisted for much of the Late Cretaceous (Engebretson et al., 1985). Convergence continued until about Mid-Eocene time, when relative motion between Kula plate (the northern part of the older Farallon plate) and North America became

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Figure 17 Location of Queen Charlotte, Georgia, and To¢no basins (adapted from Hyndman et al., 1990; Rohr and Dietricht, 1992; Katnick and Mustard, 2003). The Cretaceous portion of Queen Charlotte Basin is exposed on eastern Queen Charlotte Islands; the fault-bounded Neogene portion is mostly o¡shore. For Queen Charlotte Basin stratigraphy see Figure 16. Nanaimo Group generalized stratigraphy (Cretaceous part of Georgia Basin) is from Mustard (1994). The generalized Paleogene--Neogene subsurface stratigraphy (Point Roberts, also part of Georgia Basin) is fromYorath (1991).

increasingly strike-slip, eventually giving rise to the modern Queen Charlotte transform that separates the Paciﬁc plate from North America (from Queen Charlotte Islands north). These Early to Mid-Tertiary events coincided approximately with accretion of the Paciﬁc Rim and Crescent terranes to Wrangellia (Figure 18), and widespread volcanism on Queen Charlotte and Vancouver islands (Hyndman et al., 1990; Lewis et al., 1991). Subduction of Juan de Fuca Plate (modiﬁed Kula Plate) continued to be largely orthogonal throughout the Neogene, with igneous activity migrating eastward to the Coast belt and culminating in the modern Cascadia arc. Modern Georgia Strait and adjacent Puget Sound occupy the forearc position above the Cascadia (Juan de Fuca) subduction zone (Brandon, 2004).

7.2. Tertiary Queen Charlotte Basin The Tertiary Queen Charlotte Basin extends offshore to Hecate Strait and Queen Charlotte Sound, west to the Queen Charlotte transform plate boundary (Figures 2 and 17), and represents a signiﬁcant departure in tectonic environment compared to its Cretaceous forearc precursor. Basin subsidence was strongly inﬂuenced by rift-like stretching within an overall transform regime, where relative motion evolved progressively from transtensional to

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Figure 18 Pro¢le of the Juan de Fuca subduction zone across southernVancouver Island and Strait of Georgia, based on interpreted re£ection seismic from Hyndman et al. (1990), Cook et al. (1991), and England and Bustin (1998), showing the crustal position of Georgia and To¢no basins. CR, Crescent terrane; PR, Paci¢c Rim terrane.

pure strike-slip and ultimately transpressive (Higgs, 1991; Dehler et al., 1997; Irving et al., 2000). Thinning of the crust beneath Queen Charlotte Basin, which has accommodated up to 6 km of subsidence, has also been modeled as a lithosphere-scale, east-dipping shear zone, that was dynamically coupled to uplift of the adjacent Coast Mountains that have risen more than 3.5 km in the last 14 Myr (Rohr and Currie, 1997). Deformed Upper Cretaceous Queen Charlotte Basin strata are overlain unconformably by Paleocene to Miocene volcanic and volcaniclastic rocks, locally intercalated with non-marine sedimentary deposits. Higgs (1991) has compared these units to the syn-rift stage of the simple McKenzie (1978) stretching model. The thickest deposits are Miocene to Pliocene marine and non-marine clastics that accumulated in a complex array of extensional, or transtensional fault-bound sub-basins (Hole et al., 1993; Lyatsky and Haggart, 1993) (Figure 17). The stratigraphic architecture of the sub-basins is controlled by profound stratigraphic thickness variations that range from 200 m up to 6 km, as revealed in offshore reﬂection-seismic proﬁles (Hole et al., 1993).

7.3. Georgia (Nanaimo) Basin Georgia Basin extends across southeastern Vancouver Island, Strait of Georgia, and parts of the eastern and southern mainland (Figures 2 and 17). As such, the basin overlaps Wrangellia, the Coast belt farther east, and Cascade Mountains to the south; it postdates the Mid- to Late Cretaceous Coast Mountains thrust belt that pinned the Insular Superterrane to North America. The term Georgia Basin is used here to include Nanaimo Basin (see disputations in England and Bustin, 1998; Katnick and Mustard, 2003). Dispute also continues about the crustal position and origin of Georgia Basin. Early interpretations favor a peripheral foreland basin origin where crustal ﬂexure was driven by west-verging, Mid- to Late Cretaceous thrusting (e.g., Brandon et al., 1988; Mustard, 1994). England and Bustin (1998) on the other hand, argued that thrusting was largely over by the time Georgia Basin sedimentation had begun, and that the main locus of sedimentation did not lie immediately outboard of the thrust belt, as is usually the case in foreland-type basins. Instead, they favored a forearc setting, where the basin was located between the subduction-related trench to the west, and the Coast belt magmatic arc. The forearc interpretation may be consistent with comparable depocenters in northwest Washington (Seattle and Everett basins) that also form part of the Cascadia subduction system (Lowe et al., 2003). Basin subsidence along a northwest-trending axis began in the Turonian, and by the Maastrichtian, was sufﬁcient to accommodate more than 4 km of east-derived sediment in the Nanaimo Group (Figure 17) (Katnick and Mustard, 2003). The lower few 100 m of the Nanaimo Group contain alluvial and (economically important) coal-bearing strata that interﬁnger with shallow-marine deposits; complex facies changes were controlled primarily by paleotopography. The remainder of the succession contains submarine-fan deposits organized into laterally and vertically overlapping fan-lobe complexes, wherein inner-, mid-, and outer-fan facies are recognized (Mustard, 1994; Katnick and Mustard, 2003). Regional uplift at the end of the Cretaceous terminated deposition of the Nanaimo Group. Deformation of Nanaimo Group rocks and the underlying Wrangellian bedrock by southwest-verging imbricate thrusts resulted in 20–30% crustal shortening (Cowichan fold and thrust belt; England and Calon, 1991). Renewed subsidence in the Paleogene coincided with an eastward and southward shift in the Georgia Basin depocenter, and accumulation of 2–3 km of mainly non-marine clastics (Figure 17). Deposition continued into the Neogene with a major sedimentation focal point in the modern Fraser River delta.

7.4. Toﬁno Basin Toﬁno Basin is also linked to Cascadian subduction, but unlike modern Georgia Basin, it is located above the actively growing accretionary prism, seaward of Vancouver Island (Figures 17 and 18). The Paciﬁc Rim

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(Mesozoic) and Crescent (Eocene volcanic) terranes, that overlie the downgoing Juan de Fuca plate, were thrust beneath, and accreted to Wrangellia by Mid-Eocene (B42 Ma; Hyndman et al., 1990). These terranes formed a backstop, against which oceanic crust and sediment scraped from the subducting slab have been stacked into an accretionary prism (Figure 18). Toﬁno Basin overlies the accretionary prism and the up-dip edges of the Paciﬁc Rim and Crescent terranes, and covers most of the continental shelf west of Vancouver Island. Strata up to 4 km thick range in age from Middle Eocene to Holocene. Deformation and structural repetition of Toﬁno Basin strata are likely related to thrusting in the underlying accretionary prism, and possibly to breached compartments of overpressured ﬂuid that, in a few exploration wells, reach lithostatic pressures (Shouldice, 1971). Unconformities and complex sedimentary facies changes are a consequence of the deformation. A basin, that is broadly similar to Toﬁno Basin in age and style of deformation, is located above Yakutat terrane in the Gulf of Alaska (Figures 2 and 6). Yakutat terrane lies outboard of the northern extension of Fairweather transform, and is presently being accreted to, and partly subducted beneath southern Alaska (Bruhn et al., 2004). Neogene to Holocene sediments are accumulating above the active Yakutat accretionary prism.

7.5. Provenance linkages Nanaimo Basin has become an important target of sediment provenance investigation because of arguments concerning the hypothesis that Insular Superterrane was displaced about 3,000 km during the period 90–50 Ma (Irving and Wynne, 1991; Enkin et al., 2001; Keppie and Dostal, 2001). Deposition of the Nanaimo Group spans the ﬁrst half of this time interval and overlies Wrangellia and Coast belt rocks. Therefore, its sediment composition should reﬂect any changes in source resulting from horizontal displacement relative to the North American margin. Mahoney et al. (1999) argued that signiﬁcant populations of polycyclic Archean and Proterozoic zircons in the Nanaimo Group can only have been derived from the Canadian Shield. However, Keppie and Dostal (2001) have contested this, noting that similar zircons are found in strata of varying ages at several locations in western North America and Mexico. The issue remains contentious.

8. Discussion Sedimentary basins in the Canadian Cordillera, west of the Foreland Belt, record complex relationships between terrane accretion to ancestral North America, and the crustal–lithosphere responses (subsidence, uplift, denudation) associated with collision, subduction, rifting, and wrench tectonics. That most Cordilleran terranes traveled from afar prior to accretion is generally accepted. Some, like Stikinia, essentially traveled alone; others, like the Insular superterrane, contained lithospheric blocks that were amalgamated before docking with North America. However, vigorous debate continues concerning the distances and paleolatitudes traversed by each terrane, and to some extent the timing of terrane accretion to the North American plate margin. The attributes of each Cordilleran sedimentary basin, at scales ranging from crustal-scale basin structure, stratigraphic architecture, sedimentary facies, sediment provenance, fossils, and remnant paleomagnetic records, to rare-earth signatures in single zircon crystals, all contain the key ingredients for solving these vexing questions. All these attributes will continue to play critical roles in future investigations. The Canadian Cordillera is a vast, rugged landscape; many of its geological complexities are still to be unraveled. An optimist’s position is that there remains a great deal of exciting geoscience to be discovered.

ACKNOWLEDGMENTS Thanks to Andrew Miall for the invitation to contribute this chapter, and for persisting with the North American Phanerozoic basins project. Ray Ingersoll did an excellent job of reviewing the chapter. I also want to acknowledge the staff in the Vancouver ofﬁce of the Geological Survey of Canada, for discussions over the years, and the opportunity to work in one or two of the Cordilleran terranes. Carol Evenchick in particular provided a signiﬁcant amount of information on Bowser Basin and Stikinia.

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McKenzie, D., 1978, Some remarks on the development of sedimentary basins. Earth and Planetary Science Letters, v. 40, pp. 25–32. Mezger, J. E., Creaser, R. A., Erdmer, P., and Johnston, S. T., 2001, A Cretaceous back-arc basin in the Coast belt of the northern Canadian Cordillera: evidence from geochemical and neodymium isotope characteristics of the Kluane metamorphic assemblage, southwest Yukon. Canadian Journal of Earth Sciences, v. 38, pp. 91–103. Mihalynuk, M. G., Erdmer, P., Ghent, E. D., Cordey, F., Archibald, D. A., Friedman, R. M., and Johannson, G. G., 2004, Coherent French Range blueschist: subduction to exhumation in o2.5 m.y.? Geological Society of America Bulletin, v. 116, pp. 910–922. Monger, J. W. H., 1977, Upper Paleozoic rocks of the western Canadian Cordillera and their bearing on Cordilleran evolution. Canadian Journal of Earth Sciences, v. 14, pp. 1832–1859. Monger, J. W. H., 1989, Chapter 2. Overview of Cordilleran Geology, in Ricketts, B. D. ed., Western Canada Sedimentary Basin, Canadian Association of Petroleum Geology (Special Paper S30), pp. 9–32. Monger, J. W. H., and Ross, C. A., 1971, Distribution of fusulinaceans in the western Canadian Cordillera, Canadian Journal of Earth Sciences, v. 8, pp. 259–278. Monger, J. W. H., Price, R. A., and Tempelman-Kluit, D. J., 1982, Tectonic accretion and the origin of the two major metamorphic and plutonic welts in the Canadian Cordillera. Geology, v. 10, pp. 70–75. Monger, J. W. H., Souther, J. G., and Gabrielse, H., 1972, Evolution of the Canadian Cordillera: a plate tectonic model. American Journal of Science, v. 272, pp. 577–602. Monger, J. W. H., Wheeler, J. O., Tipper, H. W., Gabrielse, H., Harms, T., Struik, L. C., Campbell, R. B., Dodds, C. J., Gehrels, G. E., and O’Brien, J., 1991, Part B: Cordilleran Terranes; in Upper Devonian to Middle Jurassic assemblages, in Gabrielse, H. and Yorath, C. J. eds., Cordilleran Orogen in Canada, Geological Survey of Canada, Geology of Canada, no. 4, pp. 281–327. Monger, J. W. H., van der Heyden, P., Journeay, J. M., Evenchick, C. A., and Mahoney, J. B., 1994, Jurassic-Cretaceous basins along the Canadian Coast belt; their bearing on pre-Mid-Cretaceous sinistral displacements. Geology, v. 22, pp. 175–178. Murphy, D. C., van der Heyden, P., Parrish, R. R., Klepacki, D. W., McMillan, W., Struik, L. C., and Gabites, J., 1995, New geochronological constraints on Jurassic deformation of the western edge of North America, southeastern Canadian Cordillera, in Miller, D. M. and Busby, C. eds., Jurassic magmatism and tectonics of the North American Cordillera, Geological Society of America (Special Paper 299), pp. 159–171. Mustard, P. S., 1994, The Upper Cretaceous Nanaimo Group, Georgia basin, Geological Survey of Canada Bulletin 481, pp. 27–95. Osadetz, K. G., Evenchick, C. A., Ferri, F., Stasiuk, L. D., and Wilson, N. S. F., 2003, Indications for effective petroleum systems in Bowser and Sustut basins, north-central British Columbia, in Geological ﬁeldwork 2002, B.C. Ministry of Energy and Mines, Paper 2003-1, pp. 257–264. Price, R. A., 1973, Large-scale gravitational ﬂow of supracrustal rocks, southern Canadian Rockies, in de Jong, K. A. and Scholten, R. eds., Gravity and Tectonics, Wiley, New York, pp. 491–502. Ricketts, B. D., and Evenchick, C. A., 1991, Analysis of the Middle to Upper Jurassic Bowser basin, northern British Columbia, in Current Research, Part A, Geological Survey of Canada, Paper 91-1A, pp. 65–73. Ricketts, B. D., and Evenchick, C. A., 1999, Shelfbreak gullies; Products of sea-level lowstand and sediment failure: examples from Bowser basin, Northern British Columbia. Journal of Sedimentary Research, v. 69, pp. 1232–1240. Ricketts, B. D., and Evenchick, C. A., 2007, Evidence of different contractional styles along foredeep margins provided by Gilbert deltas; examples from Bowser Basin, British Columbia, Canada, Bulletin of the Canadian Society of Petroleum Geologists, v. 55, pp. 243–261. Ricketts, B. D., Evenchick, C. A., Anderson, R. G., and Murphy, D. C., 1992, Bowser basin, northern British Columbia: constraints on the timing of initial subsidence and Stikinia – North America terrane interactions. Geology, v. 20, pp. 1119–1122. Riddihough, R. R., and Hyndman, R. D., 1991, Modern plate tectonic regime of the continental margin of western Canada, in Gabrielse, H. and Yorath, C. J. eds., Geology of the Cordilleran Orogen, Geological Survey of Canada, v. 4, pp. 435–455. Riesterer, J. W., Mahoney, J. B., and Link, P. K., 2001, The conglomerate of Churn Creek: Late Cretaceous basin evolution along the Insular-Intermontane superterrane boundary, southern British Columbia. Canadian Journal of Earth Sciences, v. 38, pp. 59–73. Roddick, J. A., 1965, Vancouver North, Coquitlam and Pitt Lake map-areas, British Columbia; Geological Survey of Canada, Memoir 335. Rohr, K. M. M., and Currie, L., 1997, Queen Charlotte basin and Coast Mountains: Paired belts of subsidence and uplift caused by a lowangle normal fault. Geology, v. 25, pp. 819–822. Rohr, K. M. M., and Dietricht, J. R., 1992, Strike-slip tectonics and development of the Tertiary Queen Charlotte basin, offshore western Canada: evidence from seismic reﬂection data. Basin Research, v. 4, pp. 1–19. Rohr, K. M. M., and Furlong, K. P., 1995, Ephemeral plate tectonics at the Queen Charlotte triple junction. Geology, v. 23, pp. 1035–1038.

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Rubin, C. M., and Saleeby, J. B., 1991, The Gravina sequence: remnants of a Mid-Mesozoic oceanic arc in southern southeast Alaska. Journal of Geophysical Research, v. 96, pp. 14551–14568. Rusmore, M. E., Potter, C. J., and Umhoefer, P. J., 1988, Middle Jurassic terrane accretion along the western edge of the Intermontane superterrane, southwestern British Columbia. Geology, v. 16, pp. 891–894. Saleeby, J. B., 2000, Geochronological investigations along the Alexander-Taku terrane boundary, southern Revillagigedo Island to Cape Fox areas, southeast Alaska, in Stowell, H. H. and McClelland, W. C. eds., Tectonics of the Coast Mountains, Southeastern Alaska and British Columbia, Geological Society of America Special Paper, v. 343, pp. 107–143. Shouldice, D. H., 1971, Geology of the western Canadian continental shelf. Bulletin of Canadian Petroleum Geology, v. 19, pp. 405–436. Snyder, D. B., Clowes, R. M., Cook, F. A., Erdmer, P., Evenchick, C. A., van der Velden, A. J., and Hall, K. W., 2002, Proterozoic prism arrests suspect terranes: insights into the ancient Cordilleran margin from seismic reﬂection data. GSA Today, v. 12, pp. 4–10. Souther, J. G., 1991, Volcanic regimes, in Gabrielse, H. and Yorath, C. J. eds., Geology of the Cordilleran Orogen in Canada, Geological Survey of Canada, Geology of Canada, v. 4, pp. 457–490, Chapter 14. Souther, J. G. and Armstrong, J. E., 1966, North-central belt of the Cordillera of British Columbia, in Tectonic history and mineral deposits of the western Cordillera, Canadian Institute of Mining and Metallurgy, Special Volume 8, pp. 171–184. Stockmal, G. S., Cant, D. J., and Bell, J. S., 1992, Relationship of the stratigraphy of the Western Canada foreland basin to Cordilleran tectonics: insights from geodynamic models, in Macqueen, R. W. and Leckie, D. A. eds., Foreland Basins and Fold belts. American Association of Petroleum Geologists, Memoir 55, pp. 107–124. Struik, L. C., Schiasrizza, P., Orchard, M. J., Cordey, F., Sano, H., MacIntyre, D. G., Lapierre, H., and Tardy, M., 2001, Imbricate architecture of the Upper Paleozoic to Jurassic oceanic Cache Creek terrane, central British Columbia. Canadian Journal of Earth Sciences, v. 38, pp. 495–514. Tempelman-Kluit, D. J., 1979, Transported cataclasite, ophiolite and granodiorite in Yukon: evidence for arc-continent collision, Geological Survey of Canada, Paper 79-14, 27 pp. Thompson, R. I., Haggart, J. W., and Lewis, P. D., 1991, Late Triassic through Early Tertiary evolution of the Queen Charlotte basin, British Columbia, with a perspective on hydrocarbon potential, in Woodsworth, G. J. ed., Evolution and hydrocarbon potential of the Queen Charlotte basin, British Columbia; Geological Survey of Canada Paper 90-10, pp. 3–29. Tipper, H. W., and Richards, T. A., 1976, Jurassic stratigraphy and history of north-central British Columbia. Geological Survey of Canada, Bulletin 270, 73 pp. Tipper, H. W., Woodsworth, G. J., and Gabrielse, H., 1981, Tectonic Assemblage Map of the Canadian Cordillera. Geological Survey of Canada, Map 1505A. Trop, J. M., Ridgway, K. D., Manuszak, J. D., and Layer, P., 2002, Mesozoic sedimentary-basin development on the allochthonous Wrangellia composite terrane, Wrangell Mountains basin, Alaska: a long-term record of terrane migration and arc construction. Geological Society of America Bulletin, v. 114, pp. 693–717. Tozer, E. T., 1982, Marine Triassic faunas of North America: their signiﬁcance for assessing plate and terrane movements. Geologisch Rundschau, v. 71, pp. 1077–1104. Umhoefer, P. J., Schiarizza, P., and Robinson, M., 2002, Relay Mountain Group, Tyaughton-Methow basin, southwest British Columbia: a major Middle Jurassic to Early Cretaceous terrane overlap assemblage. Canadian Journal of Earth Sciences, v. 39, pp. 1143–1167. Unterschutz, J. L. E., Creaser, R. A., Erdmer, P., Thompson, R. I., and Daughtry, K. L., 2001, North American margin origin of Quesnellia strata in the southern Canadian Cordillera: inferences from geochemical and Nd isotopic characteristics of Triassic metasedimentary rocks. Geological Society of America Bulletin, v. 114, pp. 462–475. Van der Heyden, P., 1992, A Middle Jurassic to Early Tertiary Andean-Sierran arc model for the Coast belt of British Columbia. Tectonics, v. 11, pp. 82–97. Ward, P. L., 1995, Subduction cycles under western North America during the Mesozoic and Cenozoic eras, in Miller, D. M. and Busby, C. eds., Jurassic magmatism and tectonics of the North American Cordillera, Geological Society of America (Special Paper 299), pp. 1–45. Wheeler, J. O., and Gabrielse, H., 1972, The Cordilleran structural province, in Price, R. A. and Douglas, R. J. W. eds., Variations in tectonic styles in Canada, Geological Association of Canada (Special Paper 11), pp. 1–81. Woodsworth, G. J. and Monger, J. W. H., 1991, The Coast belt, in Yorath, C. J., 1991, Upper Jurassic to Paleogene assemblages, in Gabrielse, H. and Yorath, C.J. eds., Geology of the Cordilleran Orogen in Canada, Geological Survey of Canada, Geology of Canada, no. 4, pp. 352–354, Chapter 9. Woodsworth, G. J., Anderson, R. G., and Armstrong, R. L., 1991, Plutonic regimes, in Gabrielse, H. and Yorath, C. J. eds., Geology of the Cordilleran Orogen in Canada, Geological Survey of Canada, Geology of Canada, no. 4, pp. 491–531, Chapter 15. Yorath, C. J., 1991, Upper Jurassic to Paleogene assemblages, in Gabrielse, H. and Yorath, C. J. eds., Geology of the Cordilleran Orogen in Canada, Geological Survey of Canada, Geology of Canada, no. 4, pp. 329–371, Chapter 9. Yorath, C. J., Green, A. G., Clowes, R. M., Sutherland-Brown, A., Brandon, M. T., Kanasewich, E. R., Hyndman, R. D., and Spencer, C., 1985, Lithoprobe, southern Vancouver Island: seismic reﬂection sees through Wrangellia to the Juan de Fuca Plate. Geology, v. 13, pp. 759–762.

CHAPTER 11

Subduction-Related Sedimentary Basins of the USA Cordillera Raymond V. Ingersoll

Contents 1. Introduction 2. Devonian–Mississippian Antler Orogeny 3. Havallah Basin 4. Permo-Triassic Sonoma Orogeny 5. Post-Sonoma Successor Basin and Forearcs 6. Late Jurassic Nevadan Orogeny 7. Post-Nevadan Forearc Basins 8. Phanerozoic History 9. Conclusions Acknowledgments References

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Abstract The western USA records a complex Phanerozoic history, which includes several subduction-related orogenies with associated sedimentary basins. Older basins and the history they record are overprinted by younger events, so that much of this history must be inferred. Application of actualistic plate-tectonic models for the origin and evolution of sedimentary basins results in key constraints on interpreting the paleogeography and paleotectonics of the Cordilleran region. The history of subduction-related sedimentary basins should be viewed in the temporal and regional context of Phanerozoic evolution of the western USA, including the following phases: (1) Latest Proterozoic rifting of Rodinia to form the early Paleozoic intraplate margin, along which accumulated the Cordilleran miogeocline; (2) Devonian–Mississippian Antler orogeny and proforeland basin; (3) Mississippian-Pennsylvanian intraplate margin (including Havallah Basin), with possible transform and/or extensional tectonics; (4) Pennsylvanian-Permian Ancestral Rockies orogeny, an expression of intracontinental deformation resulting from terminal suturing of Gondwana to Laurasia along the Appalachian–Ouachita– Marathon orogenic belt; (5) Permian-Triassic Sonoma orogeny; (6)Triassic-Jurassic continental-margin magmatic arc and related basins; (7) Late Jurassic Nevadan orogeny formed during closure of a remnant ocean basin between colliding arcs; (8) Latest Jurassic-Late Cretaceous arc-trench system (including forearc basins), with major retroarc shortening and ﬂexural loading to form the Cordilleran retroforeland basin during the Late Cretaceous Sevier orogeny; (9) Latest Cretaceous-Eocene Laramide orogeny, involving shortening of continental basement far inland, as rapid convergence and possible subduction of buoyant crust resulted in ﬂat-slab subduction; (10) Oligocene ignimbrite ﬂare-up, which resulted from sudden slowing of convergence between North America and the Farallon plate, and consequent collapse of the subducting slab; (11) Miocene-Holocene triple-junction migration, with attendant formation of the Basin and Range Province, Rio Grande rift and San Andreas transform boundary. An integrated four-dimensional analysis of these diverse basins and associated terranes is necessary for a more complete understanding of the Cordillera.

1. Introduction The western USA records a complex Phanerozoic history, which includes several subduction-related orogenies with associated sedimentary basins (Figure 1). Subduction, arc magmatism, crustal collision and transform tectonics have dominated the USA Cordilleran margin or offshore since the Devonian–Mississippian Antler orogeny (e.g., Burchﬁel and Davis, 1972, 1975; Dickinson, 1976, 1981a, 1981b, 1992, 2004; Oldow et al., Sedimentary Basins of the World, Volume 5 ISSN 1874-5997, DOI 10.1016/S1874-5997(08)00011-7

r 2008 Elsevier B.V. All rights reserved.

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Figure 1 Paleozoic tectonic provinces of western USA. Also shown are signi¢cant Mesozoic and Cenozoic features (after Miller et al., 1992).

1989; Speed, 1994; Ingersoll, 1997; Stevens et al., 1997). Many types of sedimentary basins typically form during both subduction and terminal suturing of buoyant crust (e.g., Dickinson, 1974, 1992; Ingersoll, 1988; Ingersoll and Busby, 1995), including trench-slope, forearc, intraarc, backarc, retroforeland, remnant-ocean, proforeland, broken-foreland and wedge-top basins. All of these basin types have been identiﬁed in the USA Cordillera, although some types have low preservation potential and constitute a small part of the stratigraphic record (Ingersoll and Busby, 1995). This chapter emphasizes sedimentary basins formed directly by crustal collision (e.g., Antler, Sonoma and Nevadan orogenies), but other types of basins related to subduction of oceanic lithosphere are also included (e.g., forearc basins). Speciﬁcally excluded from this chapter is discussion of retroforeland, backarc and broken-foreland basins of the Ancestral Rocky Mountains (see Kluth and Coney, 1981; Kluth, 1986; Blakey, Chapter 7, this volume), Western Interior Seaway (see DeCelles, 2004; Miall et al., Chapter 9, this volume), and Laramide orogen (see Dickinson et al., 1988; Yin and Ingersoll, 1997; DeCelles, 2004; Lawton, Chapter 12, this volume). The attempted subduction of buoyant continental or magmatic-arc crust results in crustal deformation (commonly referred to as ‘‘collision’’), the creation of suture belts, and lithospheric ﬂexure to form foreland basins (Dickinson, 1974; Dewey, 1977; Ingersoll, 1988; Cloos, 1993; Speed, 1994). Primary collision-related sedimentary basins include remnant ocean basins formed between sequentially suturing crustal blocks (Graham et al., 1975), proforeland basins (peripheral foreland basins of Dickinson, 1974) formed by lithospheric ﬂexure of

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the subducting plate, and retroforeland basins (retroarc foreland basins of Dickinson, 1974) formed by lithospheric ﬂexure of the overriding plate (Dickinson, 1974; Ingersoll, 1988; Ingersoll and Busby, 1995). Subsidiary basins include wedge-top (piggyback of Ori and Friend, 1984) and backbulge (Goebel, 1991) basins related to foreland foldthrust systems (DeCelles and Giles, 1996), broken-foreland basins resulting from disruption of foreland lithosphere (commonly along preexisting crustal ﬂaws far from subduction zones and/or sutures) (Kluth, 1986; Dickinson et al., 1988; Yin and Ingersoll, 1997; Dickinson and Lawton, 2003), diverse transcurrent-fault-related basins (Christie-Blick and Biddle, 1985; Nilsen and Sylvester, 1995), impactogens (Sengor et al., 1978; Sengor, 1995) and successor basins (Ingersoll and Busby, 1995, and references therein). Remnant ocean basins are destined to be destroyed during suturing, and strata deposited in them are commonly highly deformed, and either partially subducted to form new continental crust or rapidly uplifted and eroded (Graham et al., 1975; Ingersoll et al., 1995, 2003). Thus, remnant ocean basins themselves are seldom preserved, but parts of their sedimentary ﬁll are commonly preserved as accretionary bodies (e.g., OuachitaMarathon and Songpan-Ganzi terranes) (see Miall, Chapter 8, this volume). Collisions between continents and intraoceanic magmatic arcs produce smaller accretionary bodies (e.g., modern Taiwan and Timor) that nonetheless represent the remains of signiﬁcant sedimentary basins, both pre-collisional and syncollisional. The highest ﬂux of sediment into remnant ocean basins is coincident with rapid uplift and erosion of immediately adjacent colliding crustal components (Graham et al., 1975; Ingersoll et al., 1995, 2003; Miall, Chapter 8, this volume). Deposits of remnant ocean basins commonly are the major constituents of allochthons that override continental margins during terminal suturing (e.g., Alpine Flysch units [Schwab, 1981; Homewood and Caron, 1982; Homewood, 1983; Caron et al., 1989], Apennine ﬂysch [Ricci-Lucchi, 1986] and Ouachita Flysch [Graham et al., 1975; Ingersoll et al., 1995, 2003]); these allochthons are the primary tectonic loads inducing lithospheric ﬂexure, which creates proforeland basins (Dickinson, 1974; Ingersoll and Busby, 1995; Miall, 1995). Proforeland basins are the most commonly studied collision-related basins because they form during almost all collisions and because they have relatively high preservation potential (Ingersoll and Busby, 1995). The primary example of a Cordilleran proforeland basin is the Devonian–Mississippian Antler proforeland basin of central Nevada and adjoining regions (e.g., Burchﬁel and Davis, 1972; Poole, 1974, Nilsen and Stewart, 1980; Speed and Sleep, 1982; Dickinson et al., 1983; Ingersoll, 1997). Several proforeland basins formed along the southern margin of North America during the Ouachita—Marathon–Ancestral Rockies orogeny (Kluth and Coney, 1981; Kluth, 1986; Ingersoll et al., 1995; Dickinson and Lawton, 2003), including the Fort Worth, Kerr/Val Verde and Marfa basins (Miall, Chapter 8, this volume). The broken proforeland resulting from collision of Laurasia and Gondwana included the broader Permian Basin of West Texas, reactivated Neoproterozoic failed rifts (Anadarko and Tobosa basins), and the complex uplifts and basins of the Ancestral Rocky Mountains, which are treated elsewhere in this volume (Blakey, Chapter 7, this volume; Miall, Chapter 8, this volume). Collisional retroforeland basins do not seem to have formed during the Phanerozoic in the US Cordillera, primarily because North America was part of the subducting plate in most of the collision orogenies (e.g., Antler, Ancestral Rocky and Sonoma) (Ingersoll, 1997). The only suturing that occurred while North America was the overriding plate (the Late Jurassic Nevadan orogeny) involved collision with an intraoceanic arc (Schweickert and Cowan, 1975; Ingersoll and Schweickert, 1986; Godfrey et al., 1997; Godfrey and Klemperer, 1998; Ingersoll, 2000). This complex suturing does not seem to have produced a foreland basin, although paleotectonic relations are obscured by the fact that post-Nevadan basins east of the Sierra Nevada had evolved into the Cordilleran retroforeland basins by the Cretaceous (Armstrong, 1968; Dickinson, 1976; Jordan, 1981, 1995; Ingersoll, 1997). Thus, Jurassic basins of the US Cordillera east of the magmatic arc(s) (e.g., Oldow, 1983, 1984; Oldow et al., 1989) are difﬁcult to categorize; they may include backarc (extensional or neutral settings), retroforeland and proforeland, although the latter is unlikely. Wedge-top basins and transcurrent-fault basins have low preservation potential (Ingersoll and Busby, 1995), so their absence from Paleozoic and Mesozoic orogens is to be expected. Their absence does not mean that they did not develop, only that they are not preserved. Successor basins, on the other hand, have a somewhat higher preservation potential, but their paleotectonic settings are inherently ambiguous due to their transitional nature. In one sense, all basins are successor basins because they succeed some previous paleotectonic setting (Ingersoll, 1988; Ingersoll and Busby, 1995). Overlap assemblages that date the end of orogenic or taphrogenic activity are the best candidates for the label of ‘‘successor basin’’; on the other hand, overlap assemblages should be classiﬁed according to newly developed tectonic settings wherever possible. Successor basins are characterized by the absence of regional tectonic subsidence mechanisms, such as lithospheric loading or crustal thinning (Ingersoll and Busby, 1995). Following the Late Jurassic Nevadan orogeny in northern and central California, a new subduction zone initiated within the former backarc basin of the colliding intraoceanic arc. The forearc basin thus formed could be considered a ‘‘successor basin’’ relative to the predecessor backarc basin, magmatic arc and suture zone because it was deposited as an overlap sequence that provides the minimum age for suturing of the terranes (e.g., Howell

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Figure 2 Paleolatitude trajectory and paleogeographic orientation of Cordilleran continental margin during Phanerozoic time. Shaded area indicates western USA, bounded by Canadian and Mexican borders (after Dickinson, 1992).

et al., 1985). In contrast to more typical ‘‘successor basins,’’ the Great Valley forearc basin of California and related forearc basins to the north and south provide a continuous record of arc magmatism and sedimentation spanning more than a 100 Myr. Therefore, these basins are properly classiﬁed as ‘‘forearc basins,’’ rather than as ‘‘successor basins.’’ Figure 2 shows that the USA Cordillera was on or near the equator during the early Paleozoic, with northeast trade winds dominating during mid-Paleozoic through early Mesozoic. Rapid northward movement during the Jurassic placed the USA Cordillera near its present latitude from Cretaceous to the present (see Figure 20 in Chapter 1). The following text summarizes what is known or inferred about sedimentary basins related to the major Phanerozoic orogenies of the USA Cordillera, as well as forearc and trench-slope basins deposited along the postNevadan continental margin (latest Jurassic and younger).

2. Devonian–Mississippian Antler Orogeny Early geologic mapping revealed markedly different stratigraphy and structure of eastern and western Nevada (Roberts et al., 1958; Kay and Crawford, 1964): eastern Nevada has approximately 5 km of mildly deformed, dominantly carbonate Cambrian through Mississippian strata (Miall, Chapter 5, this volume), whereas correlative strata in western Nevada consist of approximately 15 km of highly deformed clastic rocks and chert, with local volcanic and volcaniclastic intercalations. These contrasting assemblages were characterized as miogeosynclinal and eugeosynclinal, respectively, in pre-plate-tectonic terms. The recognition that western facies were thrust over eastern facies along the Roberts Mountains thrust deﬁned the Mississippian–Devonian Antler orogeny (Roberts et al., 1958; Moores, 1970; Burchﬁel and Davis, 1972, 1975; Dickinson, 1977; Nilsen and Stewart, 1980; Johnson and Pendergast, 1981; Johnson and Visconti, 1992) (Figure 1). The eugeosynclinal assemblage was transported more than 150 km eastward relative to the miogeoclinal assemblage (Harbaugh and Dickinson, 1981) within a 5–20 Myr interval spanning the Devonian–Mississippian boundary (Smith and Ketner, 1968). The Roberts Mountains allochthon includes both oceanic strata deposited far from North America and strata deposited on the continental rise and slope of the intraplate continental margin prior to accretion (Murphy et al.,

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1984). Emplacement of the Roberts Mountains allochthon was accompanied by ﬂexural downwarping of the attenuated edge of North American continental crust, with its overlying westward thickening Cambrian– Devonian shelf sequence (Stewart and Poole, 1974; Poole et al., 1992; Speed, 1994; Ingersoll, 1997; see Burgess, Chapter 2, this volume). As a result of ﬂexural loading, the Antler proforeland basin formed, and the Antler ‘‘ﬂysch’’ was deposited (Poole, 1974; Dickinson, 1977, 1981b; Speed and Sleep, 1982). Goebel (1991) suggested that the ﬂexural forebulge east of the allochthon migrated approximately 250 km eastward (based on the age of the migrating unconformity), which should be a minimum distance for migration of the allochthon. She estimated a time span of 17.5 Myr for allochthon emplacement, yielding a shortening rate of 14 mm/year. Antler foreland deposits have been recognized from north-central Idaho, across Nevada and as far south as the El Paso Mountains and Lane Mountain area of the Mojave Desert (Poole, 1974; Nilsen, 1977; Nilsen and Stewart, 1980; Greene et al., 1997; Trexler and Cashman, 1997) (Figures 1 and 3). Coeval carbonate-dominatedshelf and starved-slope deposits accumulated east of the Antler foreland basin (Gutschick et al., 1980; Sandberg and Gutschick, 1980). Similar Antler belts may be traced northward into British Columbia and possibly Yukon (Smith et al., 1993; and references therein). Complex eustatic and tectonic effects have been recognized in the Antler foreland of Montana and Idaho (e.g., Dorobek et al., 1991). The Roberts Mountains allochthon, the Antler proforeland basin and the edge of the early Paleozoic North American continental sedimentary prism can be traced from northern Nevada into east-central California (Nilsen and Stewart, 1980; Schweickert and Lahren, 1987; Greene et al., 1997; Stevens and Greene, 1999) (Figure 3). In the Inyo Mountains of southeastern California, dominantly carbonate sedimentation during the Silurian/Devonian was followed by dominantly clastic sedimentation during the Late Devonian and Early Mississippian (Stevens and Ridley, 1974). The clastic sequence includes large quartzite and carbonate clasts deposited, in part, in submarine canyons cut into a northwestward paleoslope (Stevens and Ridley, 1974). This depositional sequence may have formed during submarine uplift caused by passage of the peripheral bulge onto the edge of the continental shelf (i.e., Goebel, 1991). Remnants of the inferred intraoceanic magmatic arc and accretionary prism that overrode the edge of North America during the Antler orogeny (i.e., Speed and Sleep, 1982; Ingersoll, 1997; Dickinson, 2000) (Figure 4A–D) may be found in the metamorphic belts of the Sierra Nevada and Klamath Mountains. Detailed interpretations are controversial! The Antler and Sonoma orogenic belts can be conﬁdently traced into pendants in the eastern Sierra Nevada (Schweickert and Lahren, 1987; Greene et al., 1997) (Figures 1 and 3). Schweickert and Snyder (1981) suggested that the lower Paleozoic Shoo Fly Complex of the western metamorphic belt of the Sierra Nevada is equivalent to the Roberts Mountains allochthon. Identiﬁcation of the Antler and Sonoma belts southwest of their well-deﬁned zones in Nevada is strongly dependent on interpretations of latest Paleozoic/ earliest Mesozoic sinistral offset of the belts (i.e., Dickinson, 2000) and/or Early Cretaceous dextral offset (i.e., Schweickert and Lahren, 1990; Kistler, 1993); resolving these uncertainties is beyond the scope of this chapter. Most signiﬁcant from the perspective of the Antler proforeland model is the fact that most components of the colliding intraoceanic magmatic-arc system have been identiﬁed in the Klamath Mountains and western Sierra Nevada metamorphic belt, even if their precise paleogeographic and paleotectonic relations have not been resolved (e.g., Schweickert, 1981, 1996; Schweickert and Snyder, 1981; Hanson et al., 1988; Girty et al., 1990; Hacker and Peacock, 1990; Rubin et al., 1990; Schweickert et al., 1999; Dickinson, 2000). In the Antler proforeland basin, bathyal and neritic mudrock, sandstone, conglomerate and impure limestone (2,000–3,000 m thick) were deposited in a subsiding trough between the Roberts Mountains allochthon to the northwest (present orientation) and the neritic carbonate platform to the southeast (Poole, 1974; Nilsen and Stewart, 1980; Harbaugh and Dickinson, 1981; Goebel, 1991) (Figure 5). Poole (1974, p. 58) characterized the Antler ﬂysch as ‘‘one of the most signiﬁcant ﬂysch sequences in the western United States.’’ The term ‘‘ﬂysch’’ is used in the sense of Hsu¨ (1970) to refer to deep-marine deposits with turbidite characteristics, deposited in a ‘‘geosynclinal’’ setting (Poole, 1974). The Antler ﬂysch was deposited in the early stages of the evolution of a proforeland basin (on top of attenuated continental crust), in contrast to the generally more voluminous ﬂysch formed in remnant ocean basins (e.g., Graham et al., 1975; Ingersoll et al., 1995, 2003). The Antler ﬂysch has all the sedimentologic characteristics of typical Alpine ﬂysch (e.g., Hsu¨, 1970), including (1) hemipelagic mudrock and minor limestone, (2) turbidite sandstone, siltstone, and minor conglomerate and limestone, (3) massive disorganized conglomerate and pebbly sandstone and (4) organized conglomerate and pebbly sandstone (Poole, 1974; Nilsen, 1977; Trexler and Nitchman, 1990). Sedimentary structures include Seilacher’s (1964) Nereites ichnofacies, graded beds, basal scour, Bouma sequences, ﬂute casts, groove casts and slumps (Poole, 1974; Nilsen, 1977). Shallow-marine and nonmarine (e.g., plant) fossils were redeposited into presumed bathyal conditions in the deepest part of the foreland basin during much of its history; the foreland trough ﬁlled to above sea level near the end of its lifespan, as alluvial sediments spread eastward across the carbonate shelf (Poole, 1974; Poole and Sandberg, 1977).

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Figure 3 Generalized tectonic map showing location of northern Ritter Range pendant and elements of Antler orogenic belt from central Nevada to Mojave Desert. KF, Kern Canyon fault; MK, Mineral King pendant; SL+SM, Snow Lake and Sachse Monument pendants (after Greene et al., 1997).

Harbaugh and Dickinson (1981) documented two phases of foreland sedimentation: a retrogradational facies association of upper-slope, lower-slope, inner-fan and middle-fan facies, overlain by a progradational facies association of delta-slope, delta-front and delta-plain facies (Figures 6 and 7). These two facies associations represent early deepening and later shoaling of the foreland basin, analogous to classic Alpine ﬂysch and molasse deposition of northern Switzerland (Van Houten, 1974; Harbaugh and Dickinson, 1981). Rapid deepening

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Figure 4 True-scale actualistic analog sequential cross-sections from central California to central Utah, showing middle Paleozoic to present paleotectonic development. Mantle lithosphere (black) is de¢ned geophysically rather than geochemically. Asthenosphere is orange; rising arc magma is red. (A) Middle Devonian. (B) Late Devonian--Early Mississippian. (C) Middle Mississippian--Middle Pennsylvanian. (D) Middle Pennsylvanian--Middle Permian. (E) Late Permian--EarlyTriassic. (F) Middle Triassic--Middle Jurassic. (G) Late Jurassic. (H) Latest Jurassic. (I) Cretaceous. (J) Latest Cretaceous--Eocene. (K) Oligocene. (L) Miocene. (M) Pliocene-Quaternary (after Ingersoll, 1997).

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Figure 4 (Continued )

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Figure 4 (Continued )

Figure 5 Paleogeographic diagram showing sedimentation in Early Mississippian trough developed southeast of Antler highlands in south-central Idaho (after Nilsen, 1977).

resulted from ﬂexure of the edge of North America beneath the Roberts Mountains allochthon (Harbaugh and Dickinson, 1981; Speed and Sleep, 1982; Goebel, 1991). As shortening ended, the Roberts Mountains allochthon eroded and subsided, the foreland was ﬁlled, and the entire orogenic system was covered by the overlap assemblage of Upper Mississippian/Pennsylvanian carbonate and alluvial strata (Poole, 1974; Poole and Sandberg, 1977; Saller and Dickinson, 1982; Little, 1987; Trexler and Nitchman, 1990). Subsidence of the edge of North America to form the Antler proforeland basin is readily explained as a ﬂexural response to loading by the accretionary prism of the overriding allochthon (Speed and Sleep, 1982). A more sophisticated model is needed, however, to explain the complex erosional and sedimentologic response to ﬂexure, and the resulting stratigraphic record (e.g., Goebel, 1991; Giles and Dickinson, 1995). As the ﬂexural load migrated east/southeast, the North American shelf experienced upwarping (formation of a forebulge) east of the foredeep; beyond the forebulge would be a backbulge basin (Goebel, 1991; Giles and Dickinson, 1995; DeCelles and Giles, 1996) (Figure 8). During times of high eustatic sea level, carbonate sediments accumulated on the forebulge, whereas low eustatic sea level resulted in exposure, erosion and redistribution of sediment into

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Figure 6 Synoptic diagram illustrating key steps in evolution of Mississippian clastic sequence along northwestern £ank of Antler foreland basin: (A) Slope deposition above older shelf strata. (B) Deposition of bathyal submarine fan. (C) Deposition on prograding slope of fan delta. (D) Marine and £uvial deposition on fan-delta platform. (E) Shelf deposition. Heavy vertical line marks inferred position of measured section in central Diamond Mountains, Nevada (after Harbaugh and Dickinson, 1981).

the eastern edge of the foredeep and the western edge of the backbulge basin (Goebel, 1991; Giles and Dickinson, 1995). Eustatic sea-level changes were superposed on steady eastward migration of the forebulge (approximately 100 km by the Early Mississippian); during this eastward migration, the area previously upwarped at the forebulge subsided to form the eastern edge of the foredeep (Goebel, 1991). Emergence of the Roberts Mountains allochthon resulted in dominantly siliciclastic sedimentation in the foredeep (deposition of Chainman Shale) (Goebel, 1991). As the allochthon uplifted, sediment ﬂux to the foredeep increased; as the foredeep shoaled, the tectonic/sedimentary ﬂexural load broadened, the forebulge became more subdued, and the underﬁlled foreland evolved into an overﬁlled foreland (i.e., Jordan, 1995). The early foredeep received little clastic input, so that bedded chert and hemipelagic claystone accumulated (Pinecone sequence of the Woodruff Formation) (Goebel, 1991). The Pilot Shale and West Range Limestone accumulated in moderately deep water of the backbulge basin at the same time that shallow-water carbonate accumulated over the forebulge, except where locally emergent (Goebel, 1991). Interruption of carbonate sedimentation on the eastern Nevada shelf during latest Devonian probably signals ﬂexural upwarping (development of a forebulge of DeCelles and Giles, 1996), as the attenuated margin of North America began to feel the encroaching load of the Roberts Mountains allochthon (Goebel, 1991; Giles and

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Figure 7 Schematic cross-section of Antler foreland strata in Late Mississippian time. Horizontal dimension not to scale. Northwestern end is near longitude of Roberts Mountains; southeastern end is near longitude of Ely, Nevada. RMA is Roberts Mountains allochthon. Depiction of Antler £ysch-molasse sequence is modi¢ed from Figure 6, using same symbols (after Johnson and Visconti, 1992).

Dickinson, 1995). Mississippian isopach maps (Figure 9) show thick Lower Mississippian strata in the narrow foredeep and a broader band of thick Upper Mississippian strata, with thickest strata slightly east of Lower Mississippian depocenters (Poole, 1974). As the allochthon and foredeep migrated eastward (Poole and Sandberg, 1977; Speed and Sleep, 1982; Goebel, 1991), carbonate production kept pace with subsidence along the eastern margin of the foreland basin in eastern Nevada, where a thick Mississippian carbonate bank formed, further isolating the foredeep from cratonal sediment input. Stevens et al. (1995) attributed large-scale changes in water depth and rates of carbonate production on the shelf of southern Nevada and eastern California to ﬂexural loading, but with important eustatic inﬂuence. Rapid ﬂexural subsidence near the end of allochthon emplacement suppressed carbonate production, followed by northwestward progradation of a carbonate platform as subsidence slowed. Siliciclastic deposits of the foreland basin primarily consist of recycled-orogenic detritus derived from the allochthon (Poole, 1974; Johnson and Pendergast, 1981; Dickinson et al., 1983; Murphy et al., 1984; Cashman and Trexler, 1991; Trexler and Cashman, 1997; Dickinson and Gehrels, 2000). Speciﬁc units (e.g., Ordovician Vinini Formation) in the Roberts Mountains allochthon can be recognized as clasts within the foreland deposits (Murphy et al., 1984). The dominance of quartz (both monocrystalline and polycrystalline) and sedimentary/ metasedimentary lithic fragments, and the near absence of feldspar and volcanic lithic fragments argue strongly for derivation of the foreland deposits from an uplifted oceanic accretionary prism, analogous to modern Taiwan and Timor, related to westward (modern orientation) subduction of the edge of North America into an oceanic trench, rather than to closing of a backarc basin (Dickinson et al., 1983; Dickinson, 1985; Dorsey, 1988). Detrital zircons and sandstone petrofacies together indicate ultimate derivation of most components of the Roberts Mountains allochthon from North America (Gehrels and Dickinson, 2000; Gehrels et al., 2000), as expected during ﬁnal closure of the remnant ocean basin (e.g., Ingersoll et al., 1995, 2003). Locally emergent parts of the forebulge supplied minor detritus both eastward to the backbulge basin and westward to the foredeep (Goebel, 1991). Most of the best-studied exposures of the Roberts Mountains allochthon and foreland deposits are in central Nevada, so that general characteristics summarized herein are based primarily on observations in this area. There is suggestive evidence for northward younging of Antler deformation in Idaho (Nilsen, 1977) and southward younging of Antler orogenesis in southern Nevada and eastern California (Stevens et al., 1991; Trexler and Cashman, 1997). Trexler and Cashman (1997) demonstrated that the foredeep in southern Nevada was bathymetrically deep at the same time that the foredeep in central Nevada was ﬁlled to sea level and above (e.g., Harbaugh and Dickinson, 1981; Goebel, 1991). Much of the foredeep ﬁll in southern Nevada was derived from

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Figure 8 Early Mississippian (Kinderhook) paleogeographic setting of Antler foreland. (A) Cross-section A-Au showing position of £exural features during deposition of lower Joana Limestone. (B) Cross-section A-Au showing position of £exural features during deposition of upper Joana Limestone. (C) Map showing position of cross-section A-Au. Black dots mark locations of stratigraphic sections. Localities corrected for Cenozoic extension and Mesozoic shortening. CP, Carlin-Pinon Range; FC, Fish Creek Range; PS, Pancake Summit; RM, Roberts Mountains (after Goebel, 1991).

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Figure 9 Isopach maps (in hundreds of meters) of Mississippian strata. (A) Lower Mississippian. (B) Upper Mississippian (after Poole, 1974).

the north, presumably transported through, and derived from, shallow-marine and nonmarine depositional systems of the ﬁlled foredeep to the north (Trexler and Cashman, 1997). Persistence of deep-marine conditions to the south may reﬂect a reentrant either in the preexisting continental margin or the allochthon, or both (Trexler and Cashman, 1997). All characteristics of the Antler foreland, Roberts Mountains allochthon and related intraoceanic magmatic, sedimentary and metamorphic components of the western Sierra Nevada and Klamath Mountains are consistent with Speed and Sleep’s (1982) model for the Antler orogeny (Figure 4A–C). Reﬁnements of this model arise from study of several late Cenozoic arc-continent collisions (e.g., Taiwan, Timor and Papua New Guinea). Intraoceanic magmatic arcs involved in collisions and bordering transform margins (e.g., Lesser Antilles and Barbados) usually build accretionary prisms far more voluminous than the magmatic arcs themselves. As a result, ancient arc-continent collisions preserve accretionary prisms, but seldom preserve arcs. Most of Taiwan consists of accretionary prism; the remains of the Luzon arc are tiny by comparison. Timor is growing rapidly as northern Australia subducts beneath it, but the magmatic arc consists of small volcanic islands (e.g., Hamilton, 1979). Growth of the Apennines, the Carpathians and the Hellenic systems as accretionary belts during terminal suturing has left little evidence of related magmatic arcs (e.g., Burchﬁel and Royden, 1991), although parts of the Apennine magmatic arc remain active (e.g., Aeolian arc). Rapid slab rollback commonly accompanies these collisions (Dickinson, 2000), resulting in intraarc and backarc spreading. Following collision, backarc basins may passively subside along the new continental margin or may continue extending if other plates are involved (either or both of these scenarios probably followed the Antler orogeny; see below). Polarity reversal and initiation of a new trench commonly follows arc-continent collisions (e.g., Papua New Guinea); this scenario clearly followed the Sonoma orogeny (see below), but not the Antler orogeny. Broad proforeland basins are likely to form when a long-lived intraplate margin is involved in an arc-continent collision (e.g., modern Taiwan and Antler orogeny), whereas less well-deﬁned and preserved proforeland basins form when the colliding continent has older accreted terranes (e.g., modern Papua New Guinea and the Sonoma orogeny) (Galewsky and Silver, 1997).

3. Havallah Basin Following the Antler orogeny, uplifted parts of the arc, forearc and accretionary prism (Roberts Mountains allochthon) eroded, cooled, subsided and were buried as the new continental margin regained isostatic equilibrium (Dickinson, 1981a, 1981b, 2000; Speed and Sleep, 1982; Ingersoll, 1997) (Figure 4C and D). The eroding Roberts Mountains allochthon provided a signiﬁcant portion of the detritus to the overlap sequence and the Havallah basin (Gehrels and Dickinson, 2000; Riley et al., 2000). It has been variably proposed that the postAntler continental margin was intraplate, transform or divergent (e.g., Dickinson, 1981a, 1981b, 2000; Schweickert and Snyder, 1981; Snyder and Brueckner, 1983; Stone and Stevens, 1988a; Smith and Miller, 1990; Stevens et al., 1992; Ingersoll, 1997). In any case, sediment accumulated in the oceanic Havallah Basin west of the

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continental margin while the Antler overlap sequence accumulated on top of the Roberts Mountains allochthon; some of these strata were subsequently incorporated into the accretionary prism of Sonomia (the Golconda allochthon), prior to and during the Permo-Triassic Sonoma orogeny (Silberling and Roberts, 1962; Silberling, 1973; Speed, 1977, 1979; Miller et al., 1981, 1984; Snyder and Brueckner, 1983; Brueckner and Snyder, 1985; Stewart et al., 1986; Whiteford, 1990) (Figure 4E). The western edge of the subsiding and eroding Antler orogen was a borderland of irregular topography, amongst which the overlap sequence accumulated at the boundary between the Antler orogen and the deep-marine Havallah Basin to the west (Silberling and Roberts, 1962; Silberling, 1973; Speed, 1977; Whiteford, 1990). The Havallah Basin existed simultaneously with terminal suturing of the Ouachita-Marathon orogeny and the related Ancestral Rocky Mountains (ARM; Figure 4D) (e.g., Kluth and Coney, 1981; Kluth, 1986; Smith and Miller, 1990; Dickinson and Lawton, 2003). It is difﬁcult to evaluate in detail how events in the Ouachita— Marathon–ARM orogenic belt affected the Cordilleran margin of the USA (see Trexler et al., 1991, 2004; Blakey, Chapter 7, this volume). It is clear, however, that northeast-southwest Antler and Sonoma trends have been truncated by northwest-southeast transform boundaries, possibly during the Pennsylvanian, Permian, Triassic and/or Cretaceous at various locations (e.g., Burchﬁel and Davis, 1972, 1975; Schweickert, 1976a; Dickinson, 1981a, 1981b, 1983, 2000; Stone and Stevens, 1988a, 1988b; Schweickert and Lahren, 1990, 1991; Stevens et al., 1992; Greene et al., 1997; Dickinson and Lawton, 2001). The Havallah Basin may have originated as a backarc basin, prior to being deformed into the Golconda allochthon during the Sonoma orogeny (Burchﬁel and Davis, 1972) (Figure 4C–E). Ophiolitic crust and pelagic sediment of the Havallah Basin (e.g., Snyder and Brueckner, 1983) may have formed by backarc spreading behind (west of) the Antler arc (prior to and during the Antler orogeny), by backarc spreading behind (east of) the Sonoma arc (prior to polarity reversal to close the Havallah Basin during the Sonoma orogeny), or in some way unrelated to either arc. Overlap in ages of some pelagic strata in both the Schoonover and the Havallah sequences with ages in the Roberts Mountains allochthon and derivation of some detritus from the Roberts Mountains allochthon (Miller et al., 1981, 1984; Whiteford, 1990) suggest that Antler backarc crust underlay at least part of the Havallah Basin (Snyder and Brueckner, 1983). Extensional arc magmatism, as recorded in the northern Sierra Nevada terrane during the latest Devonian and earliest Mississippian (Harwood and Murchey, 1990), is compatible with collision of an extensional intraoceanic arc with North America during the Antler orogeny, followed by thermal subsidence and deep-marine sedimentation in the Havallah Basin. The Havallah Basin, and possible adjoining Antler remnant arc(s) would then have been incorporated into Sonomia arc basement and the Golconda allochthon during the Sonoma orogeny (e.g., Miller and Harwood, 1990; Watkins, 1990). Subsequent transform offsets and latest Jurassic Nevadan deformation have complicated this simple picture for both the Sierra Nevada and Klamath regions (Schweickert and Cowan, 1975; Ingersoll and Schweickert, 1986; Schweickert and Lahren, 1990, 1991; Schweickert et al., 1999; Dickinson, 2000; Wyld and Wright, 2001). The Havallah sequence was deposited primarily in deep water, as evidenced by the presence of radiolarian chert, distal turbidites and deep-water ichnofossils of the Nereites facies (Speed, 1977; Stewart et al., 1977; Snyder and Brueckner, 1983; Murchey, 1990; Whiteford, 1990). The presence of tholeiitic pillow basalt, hyaloclastite and pillow breccia is consistent with an oceanic-basin interpretation (Speed, 1977; Snyder and Brueckner, 1983). Over 50 percent of the Havallah sequence consists of chert and argillite, suggesting that most deposition was isolated from signiﬁcant sources of detrital sediment (Snyder and Brueckner, 1983; Murchey, 1990). The absence of thick magmatic-arc volcanic or volcaniclastic strata argues against a backarc origin for most, if not all, of the Havallah Basin (i.e., Silberling, 1975; Dickinson, 1977; Speed, 1977, 1979), although the presence of thin volcaniclastic units in the Schoonover sequence, especially in the oldest part, argues for a magmatic arc at some distance to part of the basin (Miller et al., 1981; Whiteford, 1990). Magmatic-arc, recycledorogenic and cratonal sand was contributed to the Havallah Basin by diverse sources at different times (Snyder and Brueckner, 1983; Whiteford, 1990). The variable age of pelagic sediments, and the diversity of diagenetic states during accretionary deformation together argue for prolonged down-to-the-west subduction to close the Havallah Basin, rather than short-term closing of a small basin (Snyder and Brueckner, 1983; Brueckner and Snyder, 1985; Brueckner et al., 1987) (Figure 4C–E). Complex repetition of strata of diverse ages (Upper Devonian through Permian) (Stewart et al., 1986; Babaie, 1987) also argues for a long-lived accretionary process.

4. Permo-Triassic Sonoma Orogeny During the Permian (concurrent with the end of the ARM orogeny), an intraoceanic magmatic arc (built on the microcontinent Sonomia) approached from the northwest as North American oceanic lithosphere descended below it (Moores, 1970; Speed, 1979; Dickinson, 1981a, 1981b, 2000; Ingersoll, 1997) (Figure 4D

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and E). Faunal evidence from the eastern Klamath terrane (presumed to represent part of Sonomia) (Figure 10) suggests that this terrane lay up to 6,700 km west of North America during the Early Permian (Belasky and Runnegar, 1994). In contradiction, deposition of Permian Sonomia volcanic strata on remnants of the Antler arc (presently found in the northern Sierra Nevada metamorphic belt) indicates that at least part of Sonomia was constructed near the North American margin (Dickinson, 2000; W. R. Dickinson, personal communication). In any case, by Late Permian, the intraplate or transform margin of North America had been drawn into the Sonomia subduction zone, and the Sonoma orogeny began with emplacement of the Golconda allochthon (Speed, 1977, 1979; Speed and Sleep, 1982; Snow, 1992) (Figure 4E). The Sonoma orogeny resembled the Antler orogeny in terms of overall tectonic process; however, the former differed in the following ways: (1) Sonomia was a larger landmass, (2) the North American margin was more complex prior to collision, including a residual Antler orogen, (3) a signiﬁcant proforeland basin did not form; (4) a magmatic arc is preserved, although largely buried and (5) subduction polarity reversal occurred following the Sonoma orogeny (Figure 4E and F) (Speed, 1977, 1979; Wyld, 1991; Ingersoll, 1997). Following suturing, various components of the Sonoma arc-trench system were uplifted, eroded and subsided, and/or were buried and intruded by the new continental-margin magmatic arc (Wyld, 1991, 2000; Ingersoll, 1997). There are several possible reasons why a signiﬁcant proforeland basin did not develop during the Sonoma orogeny, in contrast to during the Antler orogeny: (1) post-Antler lithosphere of the western USA was only approximately 100 Myr old during the Sonoma orogeny (difference between 350 and 250 Ma), giving it a shorter ﬂexural wavelength; (2) the continental-oceanic crustal boundary was abrupt (inherited Antler intraoceanic backarc with superposed suture complexity), in contrast to the broad attenuated early Paleozoic continental margin; and (3) inherited Antler orogenic structures produced heterogeneous lithosphere, which probably distributed stresses in complex ways. As mentioned above, Galewsky and Silver (1997) provided actualistic examples of contrasting proforeland development in young arc-continent collisions. The accreted terrane of Sonomia, which included intraarc and backarc crust west of the Golconda allochthon, consists of basaltic andesite, subordinate basaltic and siliceous rocks, along with largely volcaniclastic sedimentary rocks and less abundant carbonate, pelite and chert, representing deposition in marine platform and basinal environments (Speed, 1979; Wyld, 1990, 1991, 2000) (Figures 4D and E and 10). These rocks of Sonomia outcrop as isolated windows beneath the post-Sonoma Koipato and younger sequences (Speed, 1979). Depositional environments generally shallowed in Sonomia during the Paleozoic, with dominantly siliciclastic sedimentation during the Devono-Mississippian and dominantly intermediate-composition volcanic deposition during the Pennsylvanian (Wyld, 1990). Subsidence and sedimentation slowed in the Permian; over 3 km of Triassic arc-related sedimentary and volcanic strata then accumulated (Wyld, 1990, 1991, 2000). Faunal evidence suggests that North America and the oceanic terranes of Sonomia formed signiﬁcant distances apart (Stevens et al., 1990; Belasky and Runnegar, 1994).

5. Post-Sonoma Successor Basin and Forearcs Following the Sonoma orogeny, east-dipping subduction initiated along the Cordilleran margin (Hamilton, 1969; Dickinson, 1981a, 1981b; Ingersoll, 1997) (Figure 4E and F). The resulting continental-margin magmatic arc (Schweickert, 1976a, 1976b, 1978; Busby-Spera, 1988) obliquely overprinted older orogenic belts (i.e., Antler and Sonoma) and the edge of Precambrian crust of North America (Burchﬁel and Davis, 1972, 1975; Dickinson, 1981a, 1981b; Ingersoll, 1997). The two-dimensional model of Figure 4 is complicated by signiﬁcant truncation of the continental margin during and following the Sonoma orogeny along a sinistral transform fault (Speed, 1979; Dickinson and Lawton, 2001). Thermal subsidence of the extinct Sonoma magmatic arc and backarc, as the new continental-margin arc grew, led to deposition of marine and volcanic strata, locally over 3 km thick, from late Early Triassic through the Early Jurassic (Silberling and Roberts, 1962; Silberling, 1973; Speed, 1977, 1978; Wyld, 1990; Schweickert and Lahren, 1993; Quinn et al., 1997; Schweickert et al., 1999). These strata were deposited in the Auld Lang Syne successor basin (relative to Sonomia), which also was a backarc basin (relative to the new continental-margin magmatic arc) (e.g., Hamilton, 1969; Schweickert, 1976a, 1976b, 1978; Speed, 1978, 1979; Dickinson, 1981a, 1981b; Busby-Spera, 1988; Busby-Spera et al., 1990; Ingersoll, 1997) (Figure 4F). Shallow-marine and nonmarine conditions prevailed over wide areas of this dynamically neutral backarc region (e.g., Dickinson, 1981a, 1981b; Blakey and Gubitosa, 1983; Bilodeau and Keith, 1986; Heck and Speed, 1987; Marzolf, 1988; Riggs et al., 1996; Ingersoll, 1997). Crustal extension may have characterized some parts of the backarc (Wyld, 1990, 2000). This predominantly neutral backarc region subsequently became a retroforeland as shortening initiated during the Jurassic, ﬁrst along the Luning-Fencemaker belt, and later during the Sevier orogeny

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Figure 10 Outcrop regions of Sonomia and related terranes. Post-Sonoma orogeny rocks unpatterned. Also shown are 87Sr/86Sr contours, with .706 representing western edge of cratonal NA crust (after Speed, 1979).

(i.e., Oldow, 1978, 1983, 1984; Oldow et al., 1990; Wyld, 1996, 2002; Wyld et al., 1996; Miall et al., Chapter 9, this volume) (Figure 4G–I). The post-Sonoma convergent margin consisted of a trench, subduction complexes, forearc basins and magmatic arcs (i.e., Dickinson, 1976). Most of these early Mesozoic components are poorly understood because

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they have been destroyed, modiﬁed and/or buried during late Mesozoic and Cenozoic deformation, volcanism and sedimentation. Upper Triassic to Upper Jurassic strata of the John Day inlier of central Oregon were deposited in a forearc basin, bounded on the northwest by a subduction complex and on the southeast by the roots of a magmatic arc (Dickinson, 1979; Dickinson et al., 1979a). Little is known concerning terranes accreted northwest of the John Day inlier because they are covered, primarily by Cenozoic volcanics of the Columbia River Plateau. (See Ricketts, Chapter 10, this volume, for discussion of corresponding Canadian terranes.) Possible continuation of the John Day forearc basin and related terranes to the southwest is similarly unknown. Speed (1979) suggested that the Calaveras Complex of the Sierra Nevada foothills metamorphic belt (i.e., Schweickert et al., 1977) accreted to the southwest margin of Sonomia, although subsequent Nevadan deformation and batholithic intrusion render detailed paleogeographic reconstruction of the Calaveras Complex ambiguous. Paterson and Sample (1988), Schweickert et al. (1988, 1999) and Sharp (1988) similarly interpreted the Calaveras Complex as a post-Antler subduction complex formed along the southwestern margin of North America, which partly included accreted Sonomia. Post-Sonoma and pre-Nevadan forearc terranes of California are poorly known because they were highly deformed during the Late Jurassic Nevadan orogeny, and therefore, are best discussed in the context of the Nevadan orogeny.

6. Late Jurassic Nevadan Orogeny The Nevadan orogeny (Knopf, 1929; Taliaferro, 1942; Clark, 1964; Schweickert et al., 1984) should be viewed in the context of north-to-south suturing of diverse island arcs and related terranes to western North and South America, beginning in the Middle Jurassic in Canada (Ricketts, Chapter 10, this volume), Late Jurassic (Oxfordian) in northern California, progressing through southern California and Mexico during the Early Cretaceous, producing the Greater Antilles in the Late Cretaceous and ending in the Eocene in Columbia and Ecuador (Moores, 1998; Dickinson and Lawton, 2001). Within this general pattern is extreme complexity of timing, convergence direction, transform motion and type of collided terrane (e.g., Wilson et al., 1991). It is beyond the scope of the present discussion to attempt to resolve these complexities; rather, a brief summary of possible models for the southwestern USA is presented, with an emphasis on sedimentary basins. Conﬂicting models for the Nevadan orogeny in California were debated by Dickinson et al. (1996a, 1996b), Hopson et al. (1996) and Saleeby (1996). Ingersoll (2000, p. 395) (Figures 11–13) summarized this debate, and concluded: Four Jurassic ophiolite complexes in northern California have crystallization ages of 170–160 Ma; the Coast Range (including Great Valley) ophiolite is oldest (170–165 Ma), the Smartville ophiolite intermediate (164–160 Ma) and the Josephine ophiolite youngest (162 Ma). The Smartville ophiolite was obducted during the Sierran phase of the Nevadan orogeny (162–155 Ma), and the Josephine ophiolite was obducted during the Klamath phase of the Nevadan orogeny (153–150 Ma). The Great Valley ophiolite was not highly deformed during the Nevadan orogeny and became oceanic basement for the post-Nevadan forearc basin. Three conﬂicting models for origin of the Coast Range (including Great Valley) ophiolite have been proposed: 1. Formation by intraarc and backarc spreading related to an east-facing intraoceanic arc, which collided with a west-facing continental-margin arc during the Nevadan orogeny (Sierran phase). 2. Formation by open-ocean seaﬂoor spreading and incorporation into the continental margin during trench initiation outboard of an existing continental-margin trench. 3. Formation by forearc oblique rifting along the continental margin, followed by partial closure.

The most plausible model for formation of the Great Valley, Smartville and Josephine ophiolites is by lithospheric rifting within and/or behind magmatic arcs. I infer that the Great Valley and Smartville ophiolites formed behind an east-facing intraoceanic arc (which now constitutes most of the Foothills terrane of the Sierra Nevada metamorphic belt). The Triassic-Jurassic west-facing continental-margin forearc of California collided with the east-facing intraoceanic arc, causing the Sierran phase of the Nevadan orogeny, during which the Great Valley and Smartville ophiolites were obducted (Figure 11). Directly north of this Sierran suture zone, no collision occurred, and the continental-margin arc expanded westward, resulting in intraarc/backarc rifting to form the younger Josephine ophiolite. The Josephine ophiolite was then obducted over its arc during the Klamath phase of the Nevadan orogeny, as the Klamath trench propagated southward to trap the Great Valley ophiolite in the post-Nevadan forearc (Figure 12). One of the implications of this model is that a pre-Nevadan forearc basin and an accretionary prism (correlated with units in the Blue Mountain inlier of central Oregon) likely were partially subducted beneath the colliding intraoceanic arc and forearc during the Nevadan orogeny (Figure 11) (Schweickert, 1978; Moores and Day, 1984; Day et al., 1985; Edelman et al., 1989; Godfrey et al., 1997; Godfrey and Klemperer, 1998;

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Figure 11 True-scale actualistic analog sequential cross-sections for northern Coast Ranges, Great Valley and Sierra Nevada, consistent with map relations of Figure 12, geophysical cross-sections of Godfrey et al. (1997) and analog forward models of Ingersoll (1997). Colors and symbols same as Figure 4.Timing relations shown in cross-sections are consistent with relations along a line of latitude in the middle of the Franciscan--Great Valley--Sierra Nevada areas shown in Figure 12. Similar cross-sections, but with di¡erent ages, could be drawn along the southward propagating suture zone south of the Klamath Mountains. Abbreviations: CHGR, Copper Hill, Gopher Ridge arc complex; GVO, Great Valley ophiolite; LRPB, Logtown Ridge arc complex and 200 Myr-old Penon Blanco arc complex; SO, Smartville ophiolite (after Ingersoll, 2000).

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Figure 12 Sequential paleotectonic diagrams for Middle to Late Jurassic in northern California. Active magmatic arcs are shown with smoke, inactive without. Active subduction zones are shown by barbed symbols; suture zones are shown by suture pattern. Rifted continental margin is shown by hachured line. Active spreading centers are shown by divergent arrows on double lines, without implication of exact spreading orientation; inactive spreading centers are shown by lighter double lines without arrows. Active transform faults are indicated by thin arrows. Southward propagating trench is shown by large arrow. Abbreviations: BMF, Bear Mountain fault; CHGR, Copper Hill, Gopher Ridge arc complex; CRG, Chetco, Rogue, Galice arc complex; F, Franciscan Complex; GV, Great Valley forearc basin; GVO, Great Valley ophiolite; JO, Josephine ophiolite; LRPB, Logtown Ridge arc complex and 200 Myr-old Penon Blanco arc complex; MF, Melones fault; SF, Sonora fault; SO, Smartville ophiolite. Parentheses around ophiolite abbreviations indicate partial preservation within fault zones and/or burial beneath Great Valley forearc basin (after Ingersoll, 2000).

Schweickert et al., 1999; Ingersoll, 2000); remnants of these terranes include the Calaveras and Fiddle Creek complexes (Edelman et al, 1989; Schweickert et al., 1999). Godfrey et al. (1997) and Godfrey and Klemperer (1998) discussed geophysical evidence for these structural relations. In the southern part of the Sierra Nevada foothills metamorphic belt, all structures are west-vergent (Schweickert et al., 1988, 1999), so that the continental-margin forearc may have overridden the intraoceanic forearc in this area. In modern arc-trench systems, where two subduction complexes converge (e.g., Molucca Sea, Silver and Moore, 1978; Moore et al., 1981;

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Figure 13 Structural evolution of the Taiwan collision zone, the Ryukyu intraoceanic arc and the Okinawa Trough. The Ryukyu magmatic arc constitutes a weak zone along the continental margin. Indentation due to collision of the north end of the Luzon intraoceanic arc with the continental margin resulted in contraction in the Taiwan area; lateral extrusion north of the suture, coupled with slab rollback, formed the OkinawaTrough, initially in an intraarc setting and evolving into a backarc setting. Modi¢ed from Letouzey and Kimura (1986), based partly on Sibuet and Hsu (1997) (after Ingersoll, 2000).

Moore and Silver, 1983), extreme structural complexity results, with possible alternation of vergence in both time and space. The youngest stratigraphic unit involved in the Nevadan orogeny is the Oxfordian–Kimmeridgian Mariposa Formation (sensu lato) of the western Sierra Nevada (Imlay, 1961; Bogen, 1984). The Mariposa Formation was deposited in deep-marine environments, and was derived from mixed magmatic-arc and quartzose sources (Behrman and Parkison, 1978; Bogen, 1984; Herzig and Sharp, 1992). Structural deformation and metamorphism of the Mariposa Formation complicate paleogeographic and paleotectonic interpretations, but the most likely setting for accumulation of the Mariposa Formation was as trench-ﬁll closely associated with a western intraoceanic magmatic arc, as well as within a remnant ocean basin during north-to-south suturing of the Nevadan orogeny (Schweickert, 1978; Behrman and Parkison, 1978; Bogen, 1984; Ingersoll and Schweickert, 1986; Ingersoll, 2000) (Figures 11 and 12). The slightly younger Galice Formation in the Klamath region was deposited in a backarc basin, which opened and closed quickly during the Klamath phase of the Nevadan orogeny (Harper and Wright, 1984; Ingersoll and Schweickert, 1986; Ingersoll, 2000) (Figure 12). This tectonic model is a mirror image of the Pliocene–Holocene history of arc-continent collision to form Taiwan, which induced intraarc/backarc spreading to form the Okinawa Trough (Ingersoll, 2000; Figure 13).

7. Post-Nevadan Forearc Basins As Nevadan suturing migrated southward, oceanic lithosphere west of the Great Valley ophiolite began to subduct eastward beneath the ophiolite and associated accreted terranes (Figures 11 and 12) (Schweickert and Cowan, 1975; Ingersoll and Schweickert, 1986; Ingersoll, 2000); in the process, the Franciscan trench and subduction complex, the Great Valley forearc, and the post-Nevadan magmatic arc were born (Dickinson, 1981a, 1981b). The Great Valley forearc basin is the most thoroughly studied and best understood forearc basin on Earth, primarily because of its unique level of preservation and exposure, and easy access (Dickinson and Seely, 1979; Ingersoll, 1979, 1982a; Dickinson, 1995). In fact, it is the type forearc basin against which all other forearc basins are compared. Ojakangas (1968) was the ﬁrst to recognize and document turbidite sedimentation within the Great Valley Group, which was later classiﬁed as the ﬁll of a forearc basin by Dickinson (1970, 1971). Ojakangas (1968) documented paleocurrents and sedimentary structures typical of turbidite sedimentation in a deep-marine trough, stratigraphic trends in petrofacies reﬂecting evolutionary trends in provenance areas, and overall stratigraphic– structural relations of the basin (Ingersoll, 1999). Subsequent work (summarized by Ingersoll, 1999) built upon these aspects of basin analysis. The Great Valley forearc basin evolved from a narrow, deep, sloped basin in the latest Jurassic, through an Early Cretaceous phase dominated by basin–plain deposits, followed by Late Cretaceous diverse submarine-fan environments, and culminating in latest Cretaceous and Paleogene progradation and shoaling to sea level (Ingersoll, 1982a; Dickinson, 1995) (Figure 14). This environmental evolution occurred concurrently with

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Figure 14 Schematic cross-sections of northern California from formation of Coast Range ophiolite behind an east-facing intraoceanic arc (bottom) to termination of Great Valley forearc by conversion to transform margin (top).Vertical arrows are reference points for each cross-section and indicate location of present outcrop of Great Valley Group along west side of SacramentoValley. Stippled pattern indicates upper-slope discontinuity. TSB is trench-slope break. Foothill suture zone (BMF, MF and SF of Figure 12) indicated by diagonal and wiggly lines (from Ingersoll, 1982a).

westward and upward accretionary growth of the Franciscan Complex and retrograde migration of the Sierra Nevada magmatic arc, resulting in widening of the arc-trench gap approximately 150–80 Ma (Figures 14 and 15). Volcanic detritus from the magmatic arc dominated the Great Valley petrofacies during the Cretaceous, with subordinate metamorphiclastic and plutoniclastic input (Ojakangas, 1968; Dickinson and Rich, 1972; Ingersoll, 1978a, 1983; Mansﬁeld, 1979). The forearc basin continued to ﬁll and broaden during ﬂat-slab subduction of the Laramide orogeny (80–40 Ma), during which the magmatic arc rapidly migrated eastward (Coney and Reynolds, 1977; Dickinson and Snyder, 1978; Keith, 1978, 1982; Ingersoll, 1982a, 1997; Bird, 1984, 1988) (Figure 4H–J). Shallow-marine, deltaic and nonmarine deposits dominated as the basin ﬁlled (Cherven, 1983; Nilsen, 1990).

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Figure 15 Schematic map illustrating increase in width of Great Valley forearc basin. Approximate positions of migratory boundaries are shown for Early Cretaceous (125 Ma), mid-Cretaceous (100 Ma) and Late Cretaceous (75 Ma). Positions of western boundary at migratory trench-slope break marking inner limit of active subduction are inferred from easternmost extent of successively younger strata within Franciscan Complex. Positions of eastern boundary of Great Valley Group are primarily from subsurface information showing stratal limits. Positions of magmatic front are from western limits of radiometric dates for Sierra Nevada plutons (from Ingersoll, 1982a).

The batholithic roots of the dissected Cretaceous magmatic arc then provided most of the arkosic detritus to the uppermost Cretaceous petrofacies of the Great Valley Group (Dickinson and Rich, 1972; Ingersoll, 1978a, 1983; Mansﬁeld, 1979). As the subducted slab returned to its steeper trajectory from 40 to 20 Ma (Coney and Reynolds, 1977; Dickinson and Snyder, 1978; Keith, 1978, 1982; Ingersoll, 1997), a magmatic arc was reestablished in the Sierra Nevada; this Miocene–Pliocene magmatic arc has been short-lived, however, as northward migration of the Mendocino triple junction has converted the Franciscan subduction zone to the San Andreas transform boundary (Atwater, 1970; Dickinson and Snyder, 1979a, 1979b; Dickinson, 1981a, 1981b; Ingersoll, 1982b, 1997; Graham et al., 1983) (Figures 4K–M and 16). Neotectonic transpression along the west side of the Great Valley has locally modiﬁed structural relations among the Franciscan Complex, Great Valley ophiolite and Great Valley Group (e.g., Namson and Davis, 1988; Wentworth and Zoback., 1989; Unruh and Moores, 1992; Dickinson, 2002a). Linn et al. (1991, 1992) documented the close correspondence between geochemical characteristics of the Sierra Nevada batholith and Great Valley petrofacies. Gradual eastward migration of the magmatic arc during the Cretaceous (Figure 15) corresponded with deposition of strata with decreasing (epsilon) Nd and increasing 87 Sr/86Sr values. Linn et al. (1991, 1992) also concluded that: (1) isotopic compositions of volcanic and plutonic rocks were similar; (2) the volcanic front was denuded within a few million years and (3) most sediment was derived from the migrating volcanic front. DeGraaff-Surpless et al. (2002) and Surpless et al. (2006) reﬁned these conclusions through the study of detrital zircons from the Great Valley petrofacies. They documented abundant Nevadan and immediately postNevadan zircons (i.e., 155–145 Ma) within the Cretaceous petrofacies, results that are consistent with the development of a Nevadan suture zone overprinted by a young magmatic arc (e.g., Schweickert and Cowan, 1975; Ingersoll and Schweickert, 1986; Ingersoll, 2000) (Figures 11 and 12). As the forearc basin expanded and the magmatic arc migrated eastward during the Cretaceous, more diverse zircon populations were deposited (DeGraaff-Surpless et al., 2002). Subsidence and thermochronologic analyses of the Great Valley Group have provided additional insights. Dickinson et al. (1987) suggested that Cretaceous subsidence along the west side of the Great Valley (near the center of the Cretaceous basin) was primarily due to isostatic sediment loading on oceanic crust (Great Valley ophiolite); Moxon and Graham (1987) documented thermal subsidence along the east side of the basin during Late Cretaceous eastward migration of the magmatic arc (Ingersoll, 1979, 1982a, 1999). Both studies also documented uplift and shoaling of the basin during latest Cretaceous–Paleogene (Laramide) ﬂat-slab subduction (also see Dickinson et al., 1979b). Bostick (1974) and Dumitru (1988) documented low geothermal gradients for the Great Valley, as expected for forearcs (Ingersoll, 1999). Constenius et al. (2000) clariﬁed the complex structural development of the northwestern part of the forearc, including the history of deep-marine normal growth faults during the Early

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Figure 16 Space-time relations of major continental-margin tectonic events in California during the Miocene. Position of migrating Mendocino triple junction (northern limit of San Andreas transform system) and southern limit of Sierran arc magmatism are shown (after Graham et al., 1983).

Cretaceous; similar structural complexity may exist beneath the simple homoclinal exposure along the west side of the Great Valley. As the Great Valley forearc ﬁlled to near sea level, diverse shallow- and nonmarine environments formed; submarine canyons were cut and ﬁlled during this period, as summarized by Almgren (1978) and Dickinson et al. (1979b). Comparable forearc basins that originated south of the Great Valley are mostly dismembered by late Cenozoic deformation associated with development of the San Andreas transform boundary (Salinian block and southern California Transverse Ranges and Borderland), or they are mostly underwater along the west side of Baja California (e.g., Carey and Colburn, 1978; Link et al., 1984; Nilsen, 1986; Boehlke and Abbott, 1986; Cunningham and Abbott, 1986; Grove, 1993). Most onland remnants of the Cretaceous forearc basin west of the San Andreas fault in southern California and Salinia (e.g., Sundberg and Cooper, 1978; Nilsen and Abbott, 1981; Abbott, 1984; Link et al., 1984; Enzweiler and Bottjer, 1986; Grove, 1993) sample the youngest (eastern) parts of the forearc basin; therefore, they primarily consist of nonmarine, shallow-marine, slope and proximal-fan environments, in contrast to outer-fan and basin-plain environments represented by much of the older Sacramento Valley strata (e.g., Ingersoll, 1978b, 1979, 1982a). A few parts of the older and paleobathymetrically deeper parts of the forearc basin(s) are exposed in the Western Transverse Ranges and western Baja California (Vizcaino Peninsula and Cedros Island) (e.g., MacKinnon, 1978; Boles, 1978, 1986; Busby-Spera and Boles, 1986). Palinspastic reconstruction of southern California (e.g., Ingersoll and Rumelhart, 1999) may demonstrate an overall post-Nevadan forearc history similar to that of northern and central California, although with a younger ‘‘Nevadan orogeny’’ and subsequent forearc history (e.g., Moores, 1998; Dickinson and Lawton, 2001; Moores et al., 2002). Forearc basins north of the Great Valley (in the USA) are generally not well exposed (e.g., Dickinson, 1979; Nilsen, 1984a, 1984b; Sidle and Richers, 1985); it is unclear whether known Cretaceous strata represent several discrete basins or one or two large basins. Nilsen (1984a, 1984b, 1986) suggested that the Upper Cretaceous Hornbrook Formation, which nonconformably overlies Paleozoic rocks of the eastern Klamath Mountains, was continuous with the Ochoco basin, primarily in the subsurface, of central Oregon (Figure 17). The Hornbrook– Ochoco basin(s) was bounded on the west by accreted Klamath terranes, which provided detritus to the Hornbrook Formation (Nilsen, 1984a, 1984b), in contrast to the Great Valley forearc basin, which was primarily underlain by oceanic crust and received most of its detritus from the Sierra Nevada magmatic arc, with lesser contributions from accreted Klamath terranes (Dickinson and Rich, 1972; Ingersoll, 1978a, 1983; Mansﬁeld, 1979). Local exposures of Cretaceous strata in the coast ranges northwest of the Klamath terranes may represent trench-slope, forearc or transform borderland basins of various types (e.g., Nilsen, 1984a, 1984b; Bourgeois and Dott, 1985, 1987; Seiders and Blome, 1987). Ricketts (Chapter 10, this volume) describes coeval forearc basins in Canada. The Franciscan Complex accumulated concurrently within and adjacent to the Franciscan trench west of the Great Valley and related forearc basins (Figures 14 and 15) (Hamilton, 1969; Ernst, 1970; Dickinson, 1981a,

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Figure 17 Paleogeographic map showing inferred setting of Late Cretaceous sedimentation in northern California and Oregon (from Nilsen, 1984b).

1981b). The Franciscan Complex includes exotic components (e.g., Bailey et al., 1964; Alvarez et al., 1980; Courtillot et al., 1985; Tarduno and Alvarez, 1985; Tarduno et al., 1985; Hagstrum and Murchey, 1993), but the majority of the complex consists of locally accumulated deep-marine strata, derived predominantly from the Sierra Nevada magmatic arc and accreted terranes of the Klamath Mountains and Sierra Nevada (Dickinson et al., 1979b, 1982; Smith et al., 1979; Bachman, 1982; Seiders, 1983, 1988; Seiders and Blome, 1988). Several trenchslope basins have been identiﬁed that are structurally interweaved with more highly deformed Franciscan proper (Smith et al., 1979; Bachman, 1982).

8. Phanerozoic History The history of subduction-related sedimentary basins, as outlined above, should be viewed in the temporal and regional context of Phanerozoic evolution of the western USA (Ingersoll, 1997), including the following phases: 1. Latest Proterozoic rifting of Rodinia to form the early Paleozoic intraplate margin, along which accumulated the Cordilleran miogeocline (Stewart, 1972, 1976; Bond et al., 1985; McMenamin and McMenamin, 1990; Dalziel, 1991; Hoffman, 1991; Moores, 1991; Poole et al., 1992; Miall, Chapter 5, this volume). 2. Devonian–Mississippian Antler orogeny, as discussed above. 3. Mississippian–Pennsylvanian intraplate margin (including Havallah Basin, as discussed above), with possible transform and/or extensional tectonics (Stone and Stevens, 1988b; Dickinson and Lawton, 2001). 4. Pennsylvanian–Permian Ancestral Rockies orogeny, an expression of intracontinental deformation resulting from terminal suturing of Gondwana to Laurasia along the Appalachian–Ouachita–Marathon orogenic belt (Kluth and Coney, 1981; Kluth, 1986; Dickinson and Lawton, 2003; Blakey, Chapter 7, this volume).

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5. 6. 7. 8.

Permian–Triassic Sonoma orogeny, as discussed above. Triassic–Jurassic continental-margin magmatic arc and related basins, as discussed above. Late Jurassic Nevadan orogeny, as discussed above. Latest Jurassic–Late Cretaceous arc-trench system (including forearc basins, as discussed above), with major retroarc shortening and ﬂexural loading to form the Cordilleran retroforeland basin during the Late Cretaceous Sevier orogeny (Armstrong, 1968; Dickinson, 1976; Jordan, 1995; DeCelles, 2004; Miall, Chapter 9, this volume). 9. Latest Cretaceous–Eocene Laramide orogeny, involving shortening of continental basement far inland, as rapid convergence and possible subduction of buoyant crust resulted in ﬂat-slab subduction (Coney, 1976; Dickinson and Snyder, 1978; Livaccari et al., 1981; Bird, 1984, 1988; Engebretson et al., 1985; Cross, 1986; Spencer, 1996; DeCelles, 2004; Lawton, Chapter 12, this volume). 10. Oligocene ignimbrite ﬂare-up, which resulted from sudden slowing of convergence between North America and the Farallon plate, and consequent collapse of the subducting slab (Coney and Reynolds, 1977; Cross and Pilger, 1978; Keith, 1978, 1982; Engebretson et al., 1985; Lipman, 1992). 11. Miocene–Holocene triple-junction migration, with attendant formation of the Basin and Range Province, Rio Grande rift and San Andreas transform boundary (Atwater, 1970, 1989; Dickinson and Snyder, 1979a, b; Ingersoll, 1982b, 2001; Severinghaus and Atwater, 1990; Wernicke, 1992; Axen et al., 1993; Nicholson et al., 1994; Bohannon and Parsons, 1995; Atwater and Stock, 1998; Dickinson, 2002b).

9. Conclusions The western USA records a complex Phanerozoic history, which includes several subduction-related orogenies with associated sedimentary basins. Older basins and the history they record are overprinted by younger events, so that much of this history must be inferred. Application of actualistic plate-tectonic models for the origin and evolution of sedimentary basins results in key constraints on interpreting the paleogeography and paleotectonics of the Cordilleran region. Subduction-related sedimentation of the western USA occurred in, on and along trenches, trench slopes, forearcs, intraarcs, backarcs, retroforelands, remnant oceans, proforelands, wedge tops, broken forelands and successor basins. An integrated four-dimensional analysis of these diverse basins and associated terranes is necessary for a more complete understanding of the Cordillera. It is hoped that this synthesis of existing knowledge will help other workers in future endeavors to understand this marvelous orogenic belt.

ACKNOWLEDGMENTS I thank Andrew Miall for inviting me to write this chapter and for reviewing the manuscript. In addition, I thank Brian Ricketts for helpful comments and insights regarding Canadian basins. Kate Giles is thanked for providing Figure 8. Bill Dickinson and Rich Schweickert are also gratefully acknowledged for helpful reviews and rewarding interactions over many years.

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H., Stone, P., Dunne, G. C., Greene, D. C., Walker, J. D., and Swanson, B. J., 1997, Paleozoic and Mesozoic evolution of eastcentral California, International Geology Review, v. 39, p. 788–829. Stewart, J. H., 1972, Initial deposits in the Cordilleran geosyncline: evidence of a Late Precambrian (850 m.y.) continental separation. Geological Society of America Bulletin, v. 83, pp. 1345–1360. Stewart, J. H., 1976, Late Precambrian evolution of North America: plate tectonics implication. Geology, v. 4, pp. 11–15. Stewart, J. H., and Poole, F. G., 1974, Lower Paleozoic and uppermost Precambrian Cordilleran miogeocline, Great Basin, western United States, Society of Economic Paleontologists and Mineralogists (Special Publication 22), pp. 28–57. Stewart, J. H., MacMillan, J. R., Nichols, K. M., and Stevens, C. H., 1977, Deep-water Upper Paleozoic rocks in north-central Nevada–a study of the type area of the Havallah Formation, in Stewart, J. H., Stevens, C. H., and Fritsche, A. E. eds., Paleozoic paleogeography of the western United States, Paciﬁc Section, Society of Economic Paleontologists and Mineralogists Paciﬁc Coast Paleogeography Symposium 1, pp. 337–347. Stewart, J. H., Murchey, B., Jones, D. L., and Wardlaw, B. R., 1986, Paleontologic evidence for complex tectonic interlayering of Mississippian to Permian deep-water rocks of the Golconda allochthon in Tobin Range, north-central Nevada. Geological Society of America Bulletin, v. 97, pp. 1122–1132.

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Stone, P., and Stevens, C. H., 1988a, Pennsylvanian and Early Permian paleogeography of east-central California: implications for the shape of the continental margin and the timing of continental truncation. Geology, v. 16, pp. 330–333. Stone, P., and Stevens, C. H., 1988b, An angular unconformity in the Permian section of east-central California. Geological Society of America Bulletin, v. 100, pp. 547–551. Sundberg, F. A., and Cooper, J. D., 1978, Late Cretaceous depositional environments, northern Santa Ana Mountains, southern California, in Howell, D. G. and McDougall, K. A. eds., Mesozoic paleogeography of the western United States, Paciﬁc Section, Society of Economic Paleontologists and Mineralogists Paciﬁc Coast Paleogeography Symposium 2, pp. 535–546. Surpless, K. D., Graham, S. A., Covault, J. A., and Wooden, J. L., 2006, Does the Great Valley Group contain Jurassic strata? Reevaluation of the age and early evolution of a classic forearc basin. Geology, v. 34, pp. 21–24. Taliaferro, N. L., 1942, Geologic history and correlation of the Jurassic of southwestern Oregon and California. Geological Society of America Bulletin, v. 53, pp. 71–112. Tarduno, J. A., and Alvarez, W., 1985, Paleolatitudes of Franciscan limestones. Geology, v. 13, pp. 741. Tarduno, J. A., McWilliams, M., Debiche, M. G., Sliter, W. V., and Blake, M. C., Jr., 1985, Franciscan Complex Calera limestones: accreted remnants of Farallon plate oceanic plateaus. Nature, v. 317, pp. 345–347. Trexler, J. H., Jr., and Cashman, P. H., 1997, A southern Antler foredeep submarine fan: the Mississippian Eleana Formation, Nevada Test Site. Journal of Sedimentary Research, v. 67, pp. 1044–1059. Trexler, J. H., Jr., and Nitchman, S. P., 1990, Sequence stratigraphy and evolution of the Antler foreland basin, east-central Nevada. Geology, v. 18, pp. 422–425. Trexler, J. H., Jr., Snyder, W. S., Cashman, P. H., Gallegos, D. M., and Spinosa, C., 1991, Mississippian through Permian orogenesis in eastern Nevada: post-Antler, pre-Sonoma tectonics of the western Cordillera, in Cooper, J. D. and Stevens, C. H. eds., Paleozoic paleogeography of the western United States-II, Paciﬁc Section, Society of Economic Paleontologists and Mineralogists, Los Angeles, v. 67, pp. 317–329. Trexler, J. H., Jr., Cashman, P. H., Snyder, W. S., and Davydov, V. I., 2004, Late Paleozoic tectonism in Nevada: timing, kinematics, and tectonic signiﬁcance. Geological Society of America Bulletin, v. 116, pp. 525–538. Unruh, J. R., and Moores, E. M., 1992, Quaternary blind thrusting in the southwestern Sacramento Valley, California. Tectonics, v. 11, pp. 192–203. Van Houten, F. B., 1974, Northern Alpine molasse and similar Cenozoic sequences of southern Europe, Society of Economic Paleontologists and Mineralogists (Special Publication 19), pp. 260–273. Watkins, R., 1990, Carboniferous and Permian island-arc deposits of the eastern Klamath terrane, California, Geological Society of America (Special Paper 255), pp. 193–200. Wentworth, C. M., and Zoback, M. D., 1989, The style of Late Cenozoic deformation at the eastern front of the California Coast Ranges. Tectonics, v. 8, pp. 237–246. Wernicke, B. P., 1992, Cenozoic extensional tectonics of the U.S. Cordillera, in Burchﬁel, B. C., Lipman, P. W., and Zoback, M. L. eds., The Cordilleran orogen: coterminous U.S. [The Geology of North America, v. G-3], Geological Society of America, Boulder, pp. 553–582. Whiteford, W. B., 1990, Paleogeographic setting of the Schoonover sequence, Nevada, and implications for the Late Paleozoic margin of western North America, Geological Society of America (Special Paper 255), pp. 115–136. Wilson, K. M., Hay, W. W., and Wold, C. N., 1991, Mesozoic evolution of exotic terranes and marginal seas, western North America. Marine Geology, v. 102, pp. 311–361. Wyld, S. J., 1990, Paleozoic and Mesozoic rocks of the Pine Forest Range, northwest Nevada, and their relation to volcanic arc assemblages of the western U.S. Cordillera, Geological Society of America (Special Paper 255), pp. 219–237. Wyld, S. J., 1991, Permo-Triassic tectonism in volcanic arc sequences of the western U.S. Cordillera and implications for the Sonoma orogeny. Tectonics, v. 10, pp. 1007–1017. Wyld, S. J., 1996, Early Jurassic deformation in the Pine Forest Range, northwest Nevada, and implications for Cordilleran tectonics. Tectonics, v. 15, pp. 566–583. Wyld, S. J., 2000, Triassic evolution of the arc and backarc of northwestern Nevada, and evidence for extensional tectonism, Geological Society of America (Special Paper 347), pp. 185–207. Wyld, S. J., 2002, Structural evolution of a Mesozoic backarc fold-and-thrust belt in the U.S. Cordillera: new evidence from northern Nevada. Geological Society of America Bulletin, v. 114, pp. 1452–1468. Wyld, S. J., and Wright, J. E., 2001, New evidence for Cretaceous strike-slip faulting in the United States Cordillera and implications for terrane-displacement, deformation patterns, and plutonism. American Journal of Science, v. 301, pp. 150–181. Wyld, S. J., Quinn, M. J., and Wright, J. E., 1996, Anomalous(?) Early Jurassic deformation in the western U.S. Cordillera. Geology, v. 24, pp. 1037–1040. Yin, A., and Ingersoll, R. V., 1997, A model for evolution of Laramide axial basins in the southern Rocky Mountains, U.S.A. International Geology Review, v. 39, pp. 1113–1123.

CHAPTER 12

Laramide Sedimentary Basins Timothy F. Lawton

Contents 1. Introduction 2. Laramide Orogeny 3. Basin Distribution and Classification 4. Timing of Laramide Deformation 5. Basin Structure 6. Depositional Systems 7. Tectonics of the Laramide Orogeny Acknowledgments References

429 430 430 432 435 437 443 446 446

Abstract Sedimentary basins of the broadly defined Laramide province, which includes the Rocky Mountain region from Montana south to New Mexico, the Colorado Plateau, and the southwestern most US and northern Sonora, Mexico subsided adjacent to coeval basement-involved uplifts that began to rise as early as Turonian time and continued to develop into the Middle to Late Eocene. Basins include (1) perimeter basins, with external drainages to the Gulf of Mexico, (2) ponded basins, connected by an extensive internal drainage network for much of their histories, on the Colorado Plateau and in the central Rocky Mountains, (3) small axial basins directly east of the Colorado Plateau, and (4) inversion-flank basins adjacent to uplifted Jurassic–Early Cretaceous extensional basins south of the Colorado Plateau. Laramide basins are typically asymmetric, thickening and coarsening toward the active uplift. Deposition took place in continental settings, including alluvial fan, fluvial, lacustrine margin, and open lacustrine depositional environments. These environments shifted spatially in response to three changing factors: (1) rates of thrust faulting on the active basin margin, which determined the location of most rapid subsidence; (2) extent of the drainage networks that contributed water, solutes, and sediment to the basins, causing them to alternately spill into other basins or become isolated from one another across structural sills; and (3) climate, which also influenced water influx and freshness of lake systems. Detrital compositions in the basins generally record the rock types in nearby uplifts, with two exceptions: (1) the southern inversion-flank basins also received detritus from the Late Cretaceous magmatic arc and (2) the ponded basins of the central Rocky Mountains and northern Colorado Plateau were flooded with volcanic-lithic detritus from the Early–Middle Eocene Absaroka volcanic field. Laramide deformation above a buoyant, flat subducted slab is consistent with an observed northeastward advance of basin development during the Late Cretaceous and Paleogene.

1. Introduction Intermontane basins that developed during the Laramide orogeny in western North America succeeded the widespread Cretaceous foreland basin system linked to the Cordilleran fold and thrust belt. The Laramide basins were closely yoked to basement-cored uplifts that disrupted the former foreland basin and changed the paleogeography of the Rocky Mountain region from an extensive interior seaway to an assemblage of basins occupied by continental depositional environments that received detritus from the adjacent uplifted blocks. Depositional environments included alluvial fans, meandering and braided rivers with extensive floodplains, and lacustrine systems that varied in time and space to include lacustrine deltas, lacustrine margin shoals, intermittently exposed carbonate mudflats, open lacustrine environments, and playa settings. Basin development and basement-involved uplift began in the Late Cretaceous (as early as Turonian in northern Mexico and Campanian in the Colorado Plateau) and continued into the Middle to Late Eocene. Basement deformation Sedimentary Basins of the World, Volume 5 ISSN 1874-5997, DOI 10.1016/S1874-5997(08)00012-9

r 2008 Elsevier B.V. All rights reserved.

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Timothy F. Lawton

began earlier south of the Colorado Plateau than in the northern part of the broken foreland, and the end of deformation was earlier in the northern part of the Rocky Mountain province than in the southern part. The deformed province of Laramide basins and uplifts lay east of the former Sevier orogenic belt (Figure 1), which consisted primarily of thin-skinned thrust sheets composed mainly of Proterozoic, Paleozoic, and Mesozoic sedimentary strata, although basement was extensively involved in the hinterland of the Sevier belt (Camilleri et al., 1997). West of the thrust belt lay a region of high, altiplano-like topography that had evolved during the late development of the Sevier orogenic belt (McQuarrie and Chase, 2000; DeCelles, 2004). Laramide deformation in the central Rocky Mountain region of the United States was largely amagmatic (Dickinson and Snyder, 1978). In contrast, in the southernmost part of the Laramide deformed province, lying along the US–Mexican international boundary in southern California, Arizona, New Mexico, and Sonora, uplift and basin development were accompanied by arc magmatism that ultimately spread across the southern Laramide province (Coney and Reynolds, 1977; Damon et al., 1981; McDowell et al., 2001; Seager, 2004). Therefore, although basin structure is quite similar in the central Rocky Mountain region north and east of the Colorado Plateau (Figure 1), and the Basin-Range extensional province south of the Colorado Plateau, the influence of volcanism was much more pronounced in the south.

2. Laramide Orogeny The Laramide orogeny has been defined from both temporal and kinematic criteria, which has led to some confusion as to the meaning of the term. The event is commonly defined in a temporal context to encompass a basement-involved style of deformation that took place in the Rocky Mountain region from about 75 to 40 Ma (Coney, 1976). This definition is rendered imprecise by complete temporal overlap of thin-skinned and thickskinned deformation and by examples of basement deformation that preceded the specified time frame. Basement-involved deformation in the foreland began in Montana as early as the Early Cretaceous, prior to Late Cretaceous thin-skinned deformation (DeCelles, 1986; Schmidt et al., 1988; Perry et al., 1988). Growing evidence indicates that basement-involved deformation and basin formation in northern Sonora, southern California, and southwestern Arizona may have begun early in Late Cretaceous time (Barth et al., 2004; JacquesAyala et al., 2005), indicating substantial temporal overlap with thin-skinned thrusting in the Sevier orogenic belt. In central Utah, thin-skinned deformation near the front of the thrust belt continued until Middle to Late Eocene time and thus temporally overlapped the main phase of basement deformation that occurred farther to the east (Lawton and Trexler, 1991; Lawton et al., 1993b). Moreover, basement-involved disruption of the orogenic foreland never did take place in the Canadian Cordillera; therefore, the temporal definition of the Laramide orogeny is an artificial construct in Canada. For purposes of this chapter, the Laramide orogeny is defined on the basis of a specific basement-involved deformation style. The orogen encompasses a region of basement-cored uplifts and adjacent sediment-filled basins that formed in Late Cretaceous to Paleogene time (Figure 1). This province of uplifts and basins encircles the less-deformed Colorado Plateau on its south, east, and north sides and includes basins formed between the thrust belt and the monocline-flanked uplifts of the plateau itself. The Laramide orogen extends northward from northern Sonora, Mexico, about latitude 311 10u north, to central Montana, near latitude 461 30u north. It does not extend north into Canada, nor is it considered here to continue farther south into the Sierra Madre Oriental of Mexico, where the term Hidalogan orogeny has been proposed for coeval, largely thin-skinned deformation (Guzman and de Cserna, 1963). The best-known part of the Laramide province lies in the central Rocky Mountains of the United States, north and east of the Colorado Plateau, but roughly coeval basins of similar structural style also lie south of the Colorado Plateau, in the Basin-Range Province created by Neogene extension. These southern basins formed directly cratonward of a time-equivalent magmatic arc that migrated across the province as deformation progressed (Seager, 1983; Seager and Mack, 1986; Dickinson and Lawton, 2001; Barth et al., 2004; Seager, 2004). The southern basins are less well understood than their northern counterparts as a result of superposed younger deformation, but their presence and distribution are important to a proper understanding of the geodynamic setting of Laramide deformation and basin formation.

3. Basin Distribution and Classification Laramide basins are widely distributed on Archean crust of Wyoming and southern Montana, and they flank the relatively rigid block of the Colorado Plateau. Basins north of the Colorado Plateau and situated on Archean basement include the greater Green River, Hanna, Wind River, Laramie, Bighorn, Powder River, Crazy Mountains, Shirley, and Bull Mountain basins. Several large basins lie on the thick lithosphere of the Colorado

BM CM Archean

3a crust of

BU

Wyoming craton

Avf

0

BiU

BH

44° N

BHU

PR

OCU

300 km

WR

Late Cretaceous thrust front

WRU GMU

LaU

S

Basin Basin Types Types

ge

Perimeter

GR-GD RSU GR-W

GR-B

Ponded

an

Cr

t

us

e

ch

Ar

Ed

H L SMU

Axial Inversionflank

GR-SW

UU

NP

3b of

nd c la ni er oge t n Hi r Or ie v Se

Denver

U

t

l Be

WHU WHU

PC

FRU

F

SP S

EP WM

Colorado Plateau

CCU

HP SCU

MV

K MU

jav tra e-Ya ns va itio pa i n

KU

Ra ER NaU

SJ Yavapai crustal province (1.66-1.84Ga)

Mo

36° N

D

DCA SRU

TC

FRU

MP

40° N

l

za at az n -M ai itio ap ans v r Ya t

Mojave crustal province 1.63-1.81 Ga

SCU

G Albuquerque

B C-LJ McC

Neogene Basin-Range extensional province

RGU

3c 32° N

116° W

Tucson

Caborca block (1.69-1.78 Ga)

Gulf of California

FC

112° W

ET

LR

P

Las Cruces

Ru MZU Cab

EC

Kl

SB

CS

Mazatzal crustal province (1.63-1.75 Ga)

LHT

R-SR

108° W

ville Gren

crus

ta

e vinc l pro

-1.0

(1.3

Ga)

104° W

Figure 1 Distribution of Laramide sedimentary basins and uplifts formed between Late Cretaceous and Late Eocene time. Classi¢cation of sedimentary basins follows that of Dickinson et al. (1988) except as noted in text. Sedimentary basins: B, Baca; BH, Bighorn; BM, Bull Mountain; Cab, Cabullona; C-LJ, Carthage-La Joya; CM, Crazy Mountains; CS, Cutter Sag; D, Denver; EC, El Chanate; EP, Echo Park; ER, El Rito; F, Flagsta¡; FC, Fort Crittenden; G, Galisteo; GR, greater Green River, consisting of four subbasins (GR-B, Bridger; GR-GD, Great Divide; GR-SW, Sand Wash; GR-W,Washakie); H, Hanna; HP, Huerfano Park; K, Kaiparowits; Kl, Klondike; L, Laramie; LHT, Little Hat Top; LR, Love Ranch; McC, Upper McCoy; MP, Middle Park; MV, Monte Vista; NP, North Park; P, Potrillo; PC, Piceance Creek; PR, Powder River; Ra, Raton; R-SR, Ringbone-Skunk Ranch; Ru, Rucker; S, Shirley; SB, Sierra Blanca; SJ, San Juan; SP, South Park;TC,Table Cli¡s; U, Uinta;WR,Wind River. Basement uplifts: BHU, Black Hills; BiU, Bighorn; BU, Beartooth; CCU, Circle Cli¡s; DCA, Douglas Creek arch; FRU, Front Range; GMU, Granite Mountains; KU, Kaibab; LaU, Laramie; MZU, Montezuma; MOU, Monument; NaU, Nacimiento; OCU, Owl Creek; RSU, Rock Springs; SCU, Sange de Cristo; SMU, Sierra Madre; SRU, San Rafael; UU, Uinta;WHU,White River;WM,Wet Mountains; WRU,Wind River. Avf: Absaroka volcanic ¢eld.Thick dashed lines bound Precambrian crustal provinces of Karlstrom et al. (2004). Ages of basement provinces from Iriondo et al. (2004). Locations of structure sections in Figure 3 indicated by white circles. Sources: Dickinson et al. (1988); Fouch (1976); Goldstrand (1994); Jacques-Ayala (1999); Cather (2004); Seager (2004).

432

Timothy F. Lawton

Plateau: (1) the Uinta and Piceance Creek basins, separated by the north-trending Douglas Creek arch, lie at the northern edge of the Colorado Plateau, where they are flanked by a major reverse fault system on the southern edge of the Uinta uplift; (2) the San Juan basin, flanked by a system of monoclines on its northern and western edges and an oblique-slip reverse fault on its eastern edge (Woodward et al., 1972; Cather, 2004); and (3) the Baca basin, largely known from the subsurface at the southern extreme of the plateau. Two basins, partly dismembered by Cenozoic extension, lie along the western edge of the Colorado Plateau. Termed here the Flagstaff and Table Cliffs basins, they lay between the thrust belt and the San Rafael and Circle Cliffs uplifts, respectively (e.g., Stanley and Collinson, 1979; Goldstrand, 1994). Although depicted in Figure 1 as two discrete basins, they are separated by the Oligocene–Miocene Marysvale volcanic field and may represent a single structural basin (Hintze, 1988). The nearby Kaiparowits basin, lying between the Kaibab and Circle Cliffs uplifts and containing 800 m of Upper Campanian strata (Eaton, 1991; Roberts et al., 2005), may represent an early manifestation of Laramide deformation and deposition on the Colorado Plateau. At least nine basins lie south of the Colorado Plateau on Mesoproterozoic basement. These basins typically trend west–northwest and are elongate, with lengths of several tens of kilometers and strike-perpendicular widths of only a few kilometers to a few tens of kilometers. These include the Potrillo, Klondike, Ringbone-Skunk Ranch, Little Hat Top, Rucker, Cabullona-Fort Crittenden, El Tuli, and El Chanate basins. The Rucker and Little Hat Top basins may have once been contiguous based on their along-strike positions (e.g., Seager and Mack, 1986; Mann, 1995). The largest of these basins, and separated from the rest of the southern basins by a region of Cenozoic uplift, is the McCoy basin, a large east–west trending basin extensively dismembered by Cenozoic extension, but bordered on its north and south flanks by Late Cretaceous thrust systems that emplace crystalline rocks over the basin fill (Tosdal, 1990; Tosdal and Stone, 1994). Two major basins, the Raton and Denver basins, lie east of the Colorado Plateau, also on Mesoproterozoic basement. A string of small, narrow basins lies along the eastern edge of the Colorado Plateau (Chapin and Cather, 1981; Cather, 2004). These include the Carthage-La Joya, Galisteo, El Rito, Monte Vista, Echo Park, South Park, and North Park basins. The classification of Laramide basins remains an unsettled topic (Table 1). Chapin and Cather (1981) classified them using a system that emphasized basin structure and morphology as well as patterns of deposition within the basin. Dickinson et al. (1988) utilized a classification that stressed paleogeographic distribution of basins within the Rocky Mountain region. The basins south of the Colorado Plateau were not considered in either classification scheme. They are classified here as inversion-flank basins because they lie adjacent to and within uplifted extensional basins of Jurassic–Early Cretaceous age (Lawton, 2000; Bayona and Lawton, 2003).

4. Timing of Laramide Deformation Foreland partitioning by the rise of basement uplifts in the Rocky Mountain region generally spanned latest Cretaceous and Paleogene time, roughly 75–50 Ma. Deformation in the southern part of the Laramide province as defined here appears to have begun somewhat sooner, as early as the Turonian. Criteria for recognition of Laramide deformation vary, and include increased subsidence rates in the Late Cretaceous, onset of differential subsidence within the former Western Interior basin, reorganization of paleodispersal systems and appearance of lakes within the foreland basin system, syndepositional deformation recorded by growth strata, and apatite fissiontrack cooling ages. Because the initial rocks eroded from nascent foreland uplifts included poorly lithified shale and sandstone of the foreland basin system that was being structurally partitioned, incipient uplift is particularly difficult to recognize from changes in basin-fill petrography. Moreover, subsidence mechanisms that accommodated strata near the end of the Cretaceous remain unresolved and likely included dynamic effects related to the development of a shallow slab beneath the Laramide foreland; as a result, the initial age of Laramide deformation remains a point of disagreement. Such disagreement in turn leads to divergent viewpoints about Laramide deformational history and mechanisms for that deformation. Reports of earliest deformation are from the southern basins, where synorogenic strata derived from adjacent uplifts are generally Late Cretaceous in age, and may locally range back into the Turonian. Detrital-zircon data suggest that deposition of the El Chanate Group in northern Sonora began sometime after 107 Ma (Cenomanian) and hornblende andesite breccias interbedded with alluvial fan conglomerate of the El Tuli basin, also in northern Sonora, have recently yielded an 40Ar/39Ar laser-fusion age of 93.370.7 Ma (Turonian; Lawton et al., 2008), indicating uplift of the basin margin by that time. The Fort Crittenden Formation in southern Arizona contains Santonian–Campanian dinosaur fossils and is overlain by volcanic rocks dated near 75 Ma (Dickinson et al., 1989); the Cabullona Group in northernmost Sonora locally contains Santonian palynomorphs (Gonza´lezLeo´n and Lawton, 1995). The synorogenic Ringbone Formation of southwestern New Mexico is Late Campanian (Lucas et al., 1990; Lawton et al., 1993a) and is overlain by syndeformational andesites with latest Campanian–earliest Maastrichtian 40Ar/39Ar ages (70.5370.48 Ma, 70.6970.44 Ma, and 71.4470.38 Ma;

433

Laramide Sedimentary Basins

Table 1

Laramide basin classifications.

Basin

Dickinson et al.

Chapin & Cather

This chapter

Baca Bighorn Bull Mountain Cabullona Carthage-La Joya Crazy Mountains Cutter Sag Denver Echo Park El Chanate El Tule Flagstaff Galisteo Green River (greater) Hanna Huerfano Park Kaiparowits Klondike Laramie Little Hat Top Love Ranch Montezuma North Park-Middle Park Piceance Creek Potrillo Powder River Raton Ringbone-Skunk Ranch Rucker San Juan San Luis Shirley Sierra Blanca South Park Table Cliffs Uinta Upper McCoy Wind River

NA Ponded Perimeter NA NA Perimeter NA Perimeter NA NA NA NA Axial Ponded Axial Axial NA NA Perimeter NA NA NA Axial Ponded NA Perimeter Perimeter NA NA Perimeter NA Axial NA Axial NA Ponded NA Ponded

Green River-type Green River-type NA NA Echo Park-type NA Echo Park-type Denver-type Echo Park-type NA NA NA Echo Park-type Green River-type Green River-type Echo Park-type NA NA Green River-type NA NA NA Echo Park-type Green River-type NA Green River-type Denver-type NA NA Green River-type Echo Park-type Green River-type Denver-type Echo Park-type NA Green River-type NA Green River-type

Perimeter Ponded Perimeter Inversion-flank Axial Perimeter Perimeter Perimeter Axial Inversion-flank(?) Inversion-flank Ponded Axial Ponded Axial Axial Ponded Inversion-flank Perimeter Inversion-flank Perimeter Inversion-flank Axial Ponded Inversion-flank Perimeter Perimeter Inversion-flank Inversion-flank Perimeter Axial Axial Perimeter Axial Ponded Ponded Inversion-flank(?) Ponded

Note: (?), indicates uncertainity as to basin type.

Young et al., 2000). Similarly, andesitic breccias of the Tarahumara Formation overlying the El Chanate Group in northeastern Sonora are 71 Ma ( Jacques-Ayala, 1999). In the two previous cases, the andesites overlie synorogenic strata that record the onset of Laramide basin development, but are themselves folded and thus indicate that arc volcanism both postdated and accompanied on-going crustal deformation. Available age constraints from the upper part of the thick McCoy Mountains Formation in southern California and western Arizona also indicate Late Cretaceous deposition adjacent to basement uplifts (Tosdal and Stone, 1994; Dickinson and Lawton, 2001; Barth et al., 2004). Bracketing relationships for the age of the basin fill of the McCoy basin include detrital zircons in the upper two-thirds of the formation that range from 93 to 84 Ma (Turonian–Santonian) and postdepositional intrusion by a 73.5 Ma granodiorite (Barth et al., 2004). The detrital-zircon ages are somewhat problematic, in that they can be interpreted as either near-depositional or maximum ages, but the ages agree well with recent dates from northern Sonora. A tuff in the upper part of the McCoy Mountains Formation is 7972 Ma (U-Pb zircon Concordia intercept; Tosdal and Stone, 1994), which provides the only direct age on the formation, whose type section is about 7.5 km thick. The upper part of the McCoy Mountains Formation thus may in part predate deposition of the Fort Crittenden and Cabullona formations to the east, or it may be entirely correlative (Figure 2).

434

Timothy F. Lawton

McCoy Basin

Lutetian Danian Sel Than Campanian

Upper

Cretaceous (part)

Cenoman Turon Con San

100

Bighorn Basin

Powder River Basin

White River

Intermediate Volcanic Rocks

Duc

Duchesne River

Uin

Uinta

Bri

MB Green River Skunk Ranch

Was

Pine Hollow

Lobo Cf

San Jose

Cuchara Huerfano Farasita

Tiff Nacimiento

Torr Grand Castle

Hidalgo

tuff U-Pb McCoy Mountains (upper 3.5 km)

Laney Wilkins Pk Tipton

Colton

Poison Canyon

Puer

Pitchfork Bridger/ Washakie

Canaan Peak

Ringbone Ft. Crittenden Wahweap

Ojo Alamo

North Horn

Vermejo Trinidad

Polecat Bench

Fort Union

Wasatch

Fort Union

(U-Th)/He, Bighorn uplift

Lance

Lance

Lance Fox Hills

Incomplete

Incomplete

Incomplete

Mesaverde Group Pierre Incomplete

Menefee

Straight Cliffs

?

Wasatch

Lewis

Cabullona

youngest DZ ages

Willwood

Raton

Ojo Alamo

Kirtland/ Fruitland

Tatman

Luman

Flagstaff Mbr

Love Ranch

Kaiparowits

90

Greater Green River

Bishop Cgl

Chuska

?

80

Uinta Basin

Silicic Tuffs

Maas

70

Raton Basin

Claron

Ypresian

Eocene Paleocene

60

San Juan Basin

Green R

50

Paleogene

40

Table Cliffs Basin

Southern Basins

Chatt Bart Priab

30

Rupelian

Oligocene

Ma

Tropic

Mancos Incomplete Incomplete

Dakota

40Ar/39Ar age on tuff bed

Fluvial facies of volcaniclastic detritus

Braided fluvial facies

Lacustrine facies

Intermediate volcanic rocks

Foreland-basin strata

Eolian facies

Overbank-rich fluvial facies

Evaporitic lacustrine facies

Pediment Conglomerate

Figure 2 Correlation chart of selected Laramide basins. Formations designated as foreland basin strata underlie strata of Laramide basins, but predate foreland partitioning and formerly extended beyond bounds of Laramide structural basins. Facies indicated are dominant facies of a particular formation; most coarsen toward basin margin into conglomeratic strata, or contain intervals of other facies types. Sources: Greater Green River basin: Surdam and Stanley (1980), Machlus et al. (2002), Smith et al. (2003); Raton and San Juan basins: Cather (2004); McCoy basin:Tosdal and Stone (1994), Barth et al. (2004); Southern basins: Seager (1983), Lucas et al. (1990), Dickinson et al. (1989), Lawton et al. (1993a), Seager et al. (1997), Jacques-Ayala (1999), Dickinson and Lawton (2001);Table Cli¡s basin: Goldstrand (1994), Christensen (2005); Uinta basin: Ryder et al. (1976), Johnson (1985), Bryant et al. (1989), Machlus et al. (2002), Balls et al. (2004);Timescale: Gradstein et al. (2004).

Ages of deformation inferred from depositional patterns in the Colorado Plateau indicate onset of foreland partitioning near the end of the Campanian. In particular, the Circle Cliffs and San Rafael uplifts in Utah (Figure 1) appear to have had Late Campanian uplift histories. On the basis of 40Ar/39Ar ages on four bentonite beds, the Kaiparowits Formation in the Kaiparowits basin accumulated rapidly between 76 and 74 Ma (Late Campanian; Roberts et al., 2005). At 800 m thick, the Kaiparowits Formation is more than twice as thick as underlying Upper Cretaceous strata of the foreland basin (Eaton, 1991) in spite of its great distance from the thrust belt. It is likely that early rise of the Circle Cliffs uplift ponded, but did not completely dam, a northward draining fluvial system inferred to correlate with coeval river systems in the Book Cliffs region (Lawton, 1983, 1986; Lawton et al., 2003) to yield the observed high sediment accumulation rates of the Kaiparowits Formation. The Circle Cliffs uplift continued to pond drainage systems of the Canaan Peak Formation at the end of the Campanian (Goldstrand, 1994). A system of small inter-uplift basins persisted west of the Circle Cliffs uplift and north of the Kaibab uplift from Maastrichtian to Early Eocene time, and expanded to form the extensive palustrine–lacustrine Table Cliffs basin, recorded by the Claron Formation, in the Middle Eocene (Goldstrand, 1992, 1994). The San Rafael uplift blocked long, northeast-draining rivers and created growth strata in the upper part of the Mesaverde Group in the Late Campanian and later ponded Maastrichtian lacustrine strata between its western flank and the thrust belt (Lawton, 1983, 1986). This lacustrine system, which expanded into the

Laramide Sedimentary Basins

435

Paleogene Flagstaff and Green River lakes (Fouch, 1976), is a southwestward extension of the Uinta basin, shown as the Flagstaff basin in Figure 1. Cather (2004) interpreted a three-stage subsidence history for the San Juan and Raton basins in northern New Mexico and southern Colorado. His early phase, B80–75 Ma, is based on increased subsidence rates in stratal successions long considered as part of the Western Interior basin. The regions of increased subsidence rates, particularly in the San Juan basin, do not conform well with the long-term Laramide geometry of the basin and require changing depocenters within the basins through time (Cather, 2004). Increased rates of subsidence in the Western Interior basin are recorded by eastward expansion of isopachs in New Mexico, Utah, and Colorado, and have been attributed to the isostatic effects of a flat Farallon slab (Cross, 1986) or dynamic effects created by the decreased slab angle that preceded the flat slab and Laramide deformation (Lawton, 1994; Nummedal, 2004). Because subsidence mechanisms are difficult to assess during the Late Campanian change in geodynamic setting, determining onset of Laramide deformation from subsidence rates alone is a difficult business. Fission-track cooling ages and apatite (U-Th)/He thermochronology indicate somewhat younger Laramide uplift ages, at about 65–60 Ma in the Rocky Mountains, than are indicated by the subsidence-rate analysis. Assessment of uplift onset by these means is rendered difficult by a low Phanerozoic geothermal gradient in the Rocky Mountain region and the resultant failure of both zircon and apatite tracks to fully anneal, or basement rocks to exceed the He closure temperature, prior to Laramide uplift (e.g., Cerveny and Steidtmann, 1993; Kelley and Chapin, 1995; Crowley et al., 2002). Apatite fission-track ages suggest cooling in the Wind River Range as early as 75 Ma, with most rapid cooling between 60 and 57 Ma (Cerveny and Steidtmann, 1993). Apatite fissiontrack ages from the Front Range and Wet Mountains in Colorado suggest uplift-related cooling in the period 67–57 Ma (Bryant and Naeser, 1980; Kelley and Chapin, 1995). Similarly, apatite (U-Th)/He thermochronology suggests uplift of the Bighorn Mountains at 65 7 5 Ma (Crowley et al., 2002). The end of Laramide deformation is a source of less controversy, signaled by unconformities cutting deformed strata and generally overlain by Upper Eocene-Oligocene volcanic and volcaniclastic strata (Dickinson et al., 1988; Cather, 2004). Deformation ended near the close of the Middle Eocene in the northern part of the Rocky Mountain province, but continued to near the end of the Eocene in New Mexico and Colorado. In Montana and Wyoming, volcanic overlap successions indicate that deformation ended near the end of the Middle Eocene (Dickinson et al., 1988). Following Laramide deformation, a low relief erosion surface that truncates Laramide structures was established in Colorado and northeastern Utah. In Colorado, this surface is locally overlain by a 37–36 Ma ignimbrite (Epis and Chapin, 1975). On both flanks of the Uinta Mountains, pediments termed the Gilbert Peak erosion surface slope away from the Uinta uplift to north and south (Hansen, 1984). The surface is overlain by the Bishop Conglomerate, with a rhyolite tuff near its base that has yielded single-crystal 40Ar/39Ar ages of 34.46 7 0.26 to 33.73 7 0.11 Ma (Balls et al., 2004). In central New Mexico, the Laramide orogeny ended in Late Eocene time (B36 Ma) when shortening was followed by a shift to bimodal magmatism (Cather, 2004). Deformation in southern New Mexico is postdated by Middle Eocene andesite of the Rubio Peak Formation with a fission-track age of 41.7 Ma (Thorman and Drewes, 1980).

5. Basin Structure Laramide basins are typically paired with basement-cored uplifts that form the active basin margins. The mechanism of basement uplift, controversial for many years, is now generally considered to be crustal shortening and displacement of the uplifted block along steep to moderately dipping faults that root into the basement (Figure 3A and B; e.g., Berg, 1962; Smithson et al., 1978). The geometry of these basement faults varies significantly from the central Rocky Mountain region to the basins south of the Colorado Plateau. In the Rocky Mountain region, most basement faults dip moderately and have displacements exceeding 10 km. For example, the well imaged Wind River thrust, which emplaces crystalline basement of the Wind River Mountains over the Green River basin to the south, has 13 km of vertical separation, at least 26 km of horizontal separation and penetrates 25–30 km into the crust as a discrete fault zone (Allmendinger et al., 1985). The thrust dips 35–401 throughout its imaged length, and is planar, with no sign of steepening or flattening. Faults with similar thrust geometries have been widely drilled and imaged seismically in the Rocky Mountain region and provide abundant evidence for crustal shortening (Gries, 1983a). The location and vergence of the Laramide thrusts in the Rocky Mountain region has been attributed to reactivation of Proterozoic extensional faults (Marshak et al., 2000); however, this hypothesis can only be confirmed in the instance of the Uinta uplift, which consists of a thick Proterozoic sedimentary basin thrust both north and south over the Green River and Uinta basins, respectively (e.g., Bruhn et al., 1986). Basement-fault geometry adjacent to the southern basins is steep and thus contrasts markedly with that of the Rocky Mountain uplifts and basins. Basin-bounding faults generally trend west–northwest, dip steeply beneath

436

Timothy F. Lawton

A West Elevation, meters

East

Beartooth Uplift

5,000

Bighorn Basin

J-K C

Kl Km Kmv Kc

Tb

0

J-K

pC

Tr

J-K O-P pC

C

-5,000

B

C Blue Mountain Anticline

Uinta Basin South

North Elevation, feet

Kd Tr

5,000

B' SSW

Kf Kmv

Pw

Jn

Pp

Elevation, feet

Ju

Pm

7000

Kmc

MD C

B NNE Thunderbolt Ridge

7000

Cave Creek

lava flow

0 pC Jn

6000

Ju Tr Pw

Pp

lava flow Jg

Jg

Jce

Km

5000

MD C

Pz

6000

Jcs

Jcs

Pm

-5,000

Qal

Qal

Jct

Jct

5000

Pz

Jo Jcu?

pC 4000

4000

D Late Cretaceous Laramide basin

Late Cretaceous Laramide basin

Ku

Ku Kl

Kl

Kl

basement horse

Ju-Kl Ju-Kl

Pz

Ju-Kl

Ju-Kl Pz

Pz

null point

pC pC

T A

pC

Extensional basin formation (Late Jurassic-Early Cretaceous)

Shortening by dip-slip (Late Cretaceous)

Shortening by oblique-slip (Late Cretaceous)

Figure 3 Models for Laramide basin-margin structure. (A) Structure of east £ank of Beartooth uplift, developed during deposition of Beartooth Conglomerate (Tb) (after DeCelles et al., 1991). (B) Fold-thrust model, northern £ank of Piceance Creek basin, Colorado (redrawn from Berg, 1962). (C) Inversion structure in Chiricahua Mountains, Arizona.Thick section of Upper Jurassic strata (Jg, Glance Conglomerate, and Jct, Jcs, and Jce, members of Crystal Cave Formation) is thrust southward over Lower Cretaceous strata. Presence of Upper Jurassic rocks of Onion Saddle Formation (Jo) and undi¡erentiated Crystal Cave Formation in footwall is speculative; Lower Cretaceous strata may directly overlie Paleozoic strata, Pz.Thin horse of Paleozoic rocks along main fault is older than rocks of either footwall or hanging wall and demonstrates reactivation of Jurassic normal fault (e.g., Lawton, 2000); see model in Figure 3D (data from author’s unpublished mapping). (D) Model for active margin of inversion-£ank basins in southern New Mexico and Arizona. Laramide basins overlie thin stratigraphic sections developed during Late Jurassic--Early Cretaceous extensional basin formation, but are absent over thick sections of Upper Jurassic--Lower Cretaceous strata (after Lawton, 2000).

the uplifts, and verge to both northeast and southwest (Seager, 1983; Seager and Mack, 1986; Lawton, 2000). Where slip direction is documented, faults are oblique-slip with sinistral offset (Seager, 1983; Hodgson, 2000; Bayona and Lawton, 2003). These steep faults commonly separate Jurassic–Lower Cretaceous strata of markedly different thickness, indicating that they reactivated extensional faults of Mid-Mesozoic age and lie along the flanks of inverted sedimentary rift basins (Figure 3C and D; Lawton, 2000). Therefore, Laramide basins lie adjacent to

437

Laramide Sedimentary Basins

A SSW

Elevation, km

3

NNE

Beartooth Uplift

Northern Bighorn Basin

Tf

pC

Mz

Base Mesaverde Formation

0 Mz

Pz

pC

Pz

-3

0

20 km

B South

North Wind River Basin

Granite Mountains Twr Uplift 10,000 0 -10,000 -20,000

Owl Creek Uplift

Ti

Twr

Bighorn Basin Kml Kl

Tf pC

Tr-K

Kl

Pz

Twi

Tf Tr-K

pC Pz

Kml

Elevation, feet

0

25 miles

Figure 4 Laramide basin cross sections, showing asymmetry of ¢ll. (A) North-south section across northern part of Bighorn basin and Beartooth uplift. Tf, Fort Union Formation. Note vertical exaggeration (from Hagen et al., 1985). (B) North-south section from southern part of Bighorn basin across Owl Creek uplift toWind River basin and Granite Mountains uplift. Explanations: pC, Precambrian rocks; Pz, Paleozoic strata; Tr-K, pre-Meeteetse Mesozoic strata; Kml, Meeteetse Formation and Lewis Shale; Kl, Lance Formation; Tf, Fort Union Formation; Ti, Indian Meadows Formation; Twi,Willwood Formation; Twr,Wind River Formation (redrawn from Keefer, 1965).

the thickest parts of the former rift basins (Figure 3D), offset from the rift basins across the reactivated faults, which controlled the trend and location of the Laramide uplifts and basins. Laramide basins are asymmetric and thicken toward the active, uplifted basin margin (Figure 4). This pattern of asymmetry is present in the Rocky Mountain and southern Laramide basins, and to a lesser degree, in the basins on the east flank of the Colorado Plateau (Chapin and Cather, 1981; Yin and Ingersoll, 1997). A maximum structural relief on basement of 5–10 km is usually expressed directly adjacent to the basement uplift. Twodimensional modeling suggests that the dominant subsidence mechanism in Laramide basins was a flexural response to loading by basin-margin uplifts and the syndeformational sedimentary succession (Hagen et al., 1985). The modeled flexural profiles are consistent with bending of strong, cold lithosphere with a flexural rigidity between 1021 and 1023 Nm (Hagen et al., 1985).

6. Depositional Systems Perhaps the singular characteristic of Laramide basins that distinguishes them from the older retroarc foreland basin is their assemblage of wholly continental depositional environments. These range from alluvial fan, fluvial, deltaic, and lacustrine settings distributed according to distance from the primary sediment source of the flanking uplift. The distribution of the facies types was controlled by location of flanking uplifts and rate of deformation (Figure 5; Beck et al., 1988) as well as the position of the basin within the Laramide broken foreland (Dickinson et al., 1988). Basins distributed along the eastern flank of the Laramide province (perimeter basins of Dickinson et al., 1988) contain local lacustrine facies and fluvial facies deposited by river systems that exited the basins to the east. Conglomeratic deposits dominate basins located near the east edge of the Colorado Plateau (axial basins of Dickinson et al., 1988) due to their small size and the great structural relief of basin-margin structures (Chapin and Cather, 1981). South of the Colorado Plateau, the arc-proximal inversion-flank basins, with comparatively small lateral dimensions, likewise have basin-bounding structures with great structural relief. They contain locally derived conglomeratic alluvial deposits that interfinger with siliciclastic lacustrine strata and alternate with axial, basin-parallel fluvial systems (Hayes, 1987; Inman, 1987; Gonza´lez-Leo´n and Lawton, 1995;

438

Timothy F. Lawton

pC Pz

4 3 2 1 Mz Pz

Mudstone-coal-carbonateevaporite facies

Conglomeratic facies Sandstone-mudstone coal facies Channel sandstone bodies

pC

Mz MzPz

Pre-Upper Cretaceous strata

pC

Precambrian basement rocks

Figure 5 Generalized distribution of depositional environments with respect to basement uplift (after Beck et al., 1988). Numerals correspond to strata deposited during four phases of thrust development proposed by Beck et al. (1988): (1) Early thrusting; (2) rapid thrusting; (3) slow thrusting; and (4) post thrusting. Rapid subsidence during the ¢rst two phases results in strong asymmetric subsidence and well-de¢ned depositional axis of basin nearer the thrust front than in later two phases. Numerals are located at approximate positions of basin depositional axis.

Basabilvazo, 2000). Basins lying nearer the front of the Cordilleran thrust front (ponded basins of Dickinson et al., 1988) collected runoff from large drainage basin areas, connected to fluvial outlets via structurally elevated divides, and thus contained large freshwater to saline lakes during Middle Paleocene to Middle Eocene time when Laramide uplifts attained their greatest structural relief (Bradley, 1964; Ryder et al., 1976; Surdam and Stanley, 1979; Johnson, 1985). Depositional systems of Laramide basins fall naturally into four categories, alluvial fan, fluvial, lacustrine margin, and open lacustrine. These systems are recorded by corresponding facies associations arranged more or less concentrically about the structural basins, but strongly influenced by basin asymmetry, paleoslope orientation and direction and distance to source uplifts. The alluvial fan system represents the most source-proximal deposits of Laramide intermontane basins. Strata consist of massive, crudely laminated pebble to boulder conglomerate and subordinate sandstone deposited by debris-flow and sheet-flood processes (Andersen and Picard, 1974; Ryder et al., 1976). These deposits are typically restricted to regions that lay within a few kilometers of active uplifts (Steidtmann, 1971; Surdam and Stanley, 1980), and commonly preserve internal angular unconformities that record uplift of the adjacent range (Bryant et al., 1989; DeCelles et al., 1991). Clast compositions match the range of rock types present in nearby uplifts and have been used to discriminate individual local uplifts that flanked Laramide basins (e.g., Steidtmann, 1971). Conglomeratic successions commonly contain inverted-clast stratigraphies that record progressive erosion through the sedimentary cover and into the basement of the uplifts (Hansen, 1984; Clemons and Mack, 1988; DeCelles et al., 1991; Cather, 2004). Conglomerate and sandstone in inversion-flank basins contain abundant clasts and grains derived from Upper Jurassic and Lower Cretaceous rocks of the older rift-basin strata that were uplifted during basin inversion (Gonza´lez-Leo´n and Lawton, 1995; Mann, 1995; Basabilvazo, 2000; Dickinson and Lawton, 2001). The fluvial system consists of channel-belt and floodplain environments (e.g., Ryder et al., 1976; Bown and Kraus, 1987; Kraus, 1987). Fluvial systems of Laramide basins lay downstream of and adjacent to alluvial fan environments and upstream of the lacustrine margin, or, if lakes were not developed in the basin, formed axial drainages parallel to the structural axis of the basin (Figure 5; e.g., Steidtmann, 1971; Beck et al., 1988). Fluvial systems record alluvial plains traversed by braided or meandering channels, and overbank settings with extensive paleosols are common in basins with hydrologic outlets (e.g., Willwood Formation of Bighorn basin; Bown and Kraus, 1987; Kraus, 1987). In basins with well-developed lacustrine systems, alluvial plains varied tremendously

439

Laramide Sedimentary Basins

in length. Adjacent to active uplifts, the alluvial plains formed short links between the alluvial fan system and lacustrine margins. Examples of short fluvial systems are recorded in the Lower Eocene Wasatch Formation, which was derived from the north flank of the Uinta uplift adjacent to the Sand Wash subbasin (Surdam and Stanley, 1980) and the coarse-grained Duchesne River Formation, which was derived from sedimentary rocks of the Uinta uplift directly to the north of the Uinta basin (Andersen and Picard, 1974; Dickinson et al., 1986). Alluvial plains on the distal flanks of basins in places stretched for hundreds of kilometers and have been reconstructed from their distinctive detrital compositions. Fluvial dispersal systems transported detritus into the Uinta and Piceance basins from flanking highlands that lay different distances from the basin. The North Horn Formation, which includes local lacustrine strata that record initial ponding of eastbound thrust belt-derived drainage systems by the San Rafael uplift, also contains detritus derived from continuing uplift of the Sevier orogenic belt (Lawton et al., 1993b). The Colton Formation, consisting of arkosic detritus that interfingers with the Green River Formation on the south flank of the Uinta basin, was derived from basement uplifts in southcentral Colorado, located approximately 250 km from the basin (Dickinson et al., 1986). The younger Middle Eocene Crazy Hollow Formation overlies the Green River Formation of the Flagstaff basin and records delivery of chert-lithic detritus to the southwestern arm of Lake Uinta from southwestern Utah, northwestern Arizona, and southern Nevada (Weiss and Warner, 2001). Fluvial sandstone compositions are typically lithic-rich, with the lithic grain types varying widely depending upon the composition of the rock types in nearby uplifts. Quartz-lithic sandstones with abundant sedimentary grains are common in basins adjacent to uplifts, such as the Uinta uplift, whose sedimentary cover was never completely removed (Isby and Picard, 1985; Dickinson et al., 1986). Volcanic-lithic grains derived through longdistance transport of detritus from the Absaroka volcanic field are present in Eocene strata of the Green River, Piceance Creek, and Uinta basins (Surdam and Stanley, 1980; Johnson, 1985). Volcanic-lithic detritus is also present, in deposits of axial drainages of the inversion-flank basins in the southern part of the Laramide province (Inman, 1987; Basabilvazo, 2000; Dickinson and Lawton, 2001; Barth et al., 2004). The volcanic grains were derived from older volcanic rocks of Jurassic age in the adjoining uplifts and from time-equivalent volcanism in the Cordilleran arc. Detrital-zircon data from the McCoy Mountains Formation augment compositional data and offer new insight into the nature of basin-margin sources of the McCoy basin. Detrital-zircon age probability plots from the Late Cretaceous part of the formation that is unambiguously tied to basement-involved deformation (Figure 6; Barth et al., 2004) are distinctively bimodal, with Jurassic–Late Cretaceous grain ages representing detritus from the volcanic arc to the west and southwest, and a population near 1,700 Ma derived from uplifted Proterozoic basement of the basin margins (Figure 1). The lacustrine margin system consists of an assemblage of siliciclastic and carbonate strata that record deltaic, interdeltaic, and littoral environments. These environments shifted laterally tens of kilometers as lakes within the

84

Sample MC5 n=32 zircon grains

Relative probability

Grains derived from magmatic arc sources to west and southwest 162 Grains derived from local Proterozoic basement of basin-margin uplift

0

500

1000

1500

1682

2000

Siltstone Member, Age in Ma

Figure 6 Detrital-zircon age probability plot for sample from uppermost part of McCoy Mountains Formation reported by Barth et al. (2004). Abscissa is age in millions of years; peaks represent relative probability of encountering a particular grain age in the Zircon population. Numbers near peaks are peak ages in millions of years. Youngest ages in this sample of 84 Ma indicate maximum depositional age of upper part of McCoy Mountains Formation.

440

Timothy F. Lawton

basins expanded and contracted, resulting in complexly interfingering lithofacies associations (Ryder et al., 1976; Surdam and Stanley, 1979; Johnson, 1985). Siliciclastic rocks include gray-green calcareous claystone with calcareous nodules and channelform sandstone that accumulated in deltaic and interdeltaic settings. Carbonate strata include grain- and mud-supported rocks deposited on lake-margin strandlines and carbonate flats, respectively. Dolomicrite strata in the Green River Formation were formed on exposed carbonate mudflats that were subject to evaporative pumping during lake lowstands (Surdam and Stanley, 1979). Upward-coarsening siliciclastic cycles interpreted as deltaic successions indicate that lake-margin water depths in the greater Green River basin were on the order of 2 m, but uncommon Gilbert delta foresets indicate local water depths of 25 m on the north side of the basin (Surdam and Stanley, 1979). The open lacustrine depositional system is represented by carbonate mudstone, commonly varved and kerogenous, fossiliferous carbonate wackestone, and claystone with minor sandstone and siltstone. Strata of the open lacustrine system are generally assigned to the Green River Formation in the Flagstaff, Uinta, Piceance Creek, and greater Green River basins. The varved kerogenous carbonate mudstone, also termed oil shale, was deposited in lake-center environments removed from sources of detrital influx. Evaporitic strata, including bedded trona and halite with interbedded oil shale, are present at several levels of the Green River Formation in the greater Green River basin (Eugster and Hardie, 1975; Surdam and Stanley, 1979; Smoot, 1983; Sullivan, 1985; Pietras et al., 2003b). Nahcolite and halite are interbedded with oil shale in the Uinta basin ( Johnson, 1985). Saline conditions are also indicated in carbonate mudstone by the presence of saline mineral casts and molds in some horizons and lack of freshwater mollusks (Surdam and Stanley, 1979; Johnson, 1985). Nearshore environments contain abundant mollusks and ostracodes (Ryder et al., 1976; Johnson, 1985). Cycles of kerogenrich laminated carbonate mudstone bearing evidence for exposure and shallow water conditions, such as desiccation polygons and flat-pebble intraclast conglomerate, and structureless dolomicrite beds averaging 2 m thick are present in the greater Green River basin (Surdam and Stanley, 1980). The open lacustrine system has been interpreted alternatively as a perennially deep, salinity-stratified lake (Bradley and Eugster, 1969), or a playa setting (Eugster and Surdam, 1973; Eugster and Hardie, 1975). The former interpretation explains the varved carbonate mudstone, whereas the latter explains intervals of evaporitic strata with desiccation features. To reconcile these end-member hypotheses, Surdam and Stanley (1979) invoked water depth changes above a nearly flat lake bottom over which slight changes in water depth resulted in widespread transgressions and regressions. Kerogenous laminated carbonate accumulated as a result of lakebottom algal productivity (Eugster and Surdam, 1973) during maximum lake highstands, although water depths may not have exceeded 2 m. During lowstands, wide exposed carbonate mudflats were altered to dolomicrite by evaporative pumping. Current models of deposition in the greater Green River basin favor alternations from fluvial-lacustrine conditions deposited during times when influx of water and sediment fill exceeded potential accommodation (overfilled lake model) through a fluctuating profundal facies association of kerogenous carbonate mudstone when sediment fill and water influx roughly balanced accommodation (balanced-fill lake model) to an evaporative facies association of linked mudflat–playa settings recorded by evaporite-bearing strata when rates of potential accommodation exceeded sediment fill and water influx (underfilled lake model; Carroll and Bohacs, 1999). Interaction of Laramide depositional systems and the influence upon them of basin hydrology, tectonics, and long-term climate are illustrated by the Lower–Middle Eocene Green River Formation and coeval flanking fluvial wedges of the interior ponded basins (Figure 7). These basins include the greater Green River basin, divided into the Bridger, Great Divide, Washakie and Sand Wash subbasins, and the Piceance Creek and Uinta basins. The fill of the greater Green River basin thickens westward toward the Wyoming salient of the thrust belt and southward toward the Uinta uplift, both of which were important sediment sources (Sullivan, 1985; Surdam and Stanley, 1979). The north-trending Rock Springs uplift of latest Cretaceous age (Gries, 1983b) partitions the greater Green River basin. Other Laramide sediment sources included the Granite Mountains and Sierra Madre uplifts on the east side of the basin, the Wind River uplift on the north flank, and beginning in Middle Eocene time, the Absaroka volcanic field lying north of the Wind River uplift (Figure 8). The Uinta and Piceance Creek basins thicken northward toward the Uinta uplift and are separated by the north-trending Douglas Creek arch, which influenced stratigraphic thickness trends into the Early Eocene ( Johnson, 1985). In the Paleocene, the Uinta and Piceance Creek basins received basement detritus from the south, and the Uinta basin also received sediment from the thrust belt. In the Eocene, the Uinta Mountains became the primary source for clastic detritus in the Uinta basin (Andersen and Picard, 1974; Ryder et al., 1976; Dickinson et al., 1986). The Green River Formation represents deposits of Lake Gosiute and Lake Uinta north and south of the Uinta uplift, respectively (Figure 8). In the greater Green River basin, the Green River Formation is divided into four members, the Luman Tongue and Tipton, Wilkins Peak and Laney members (Figure 7; Sullivan, 1985; Carroll and Bohacs, 1999; Smith et al., 2003). These members are lacustrine margin and open lacustrine deposits; the Luman and Tipton members record freshwater fluvio-lacustrine conditions that changed near the termination of

441

Laramide Sedimentary Basins

West

Bridger subbasin Bridger Formation (volcanic lithic)

Rock Springs uplift

Sand Wash subbasin

Washakie Formation (volcanic lithic)

46.83 ± 0.90 47.56 ± 0.14 48.65 ± 0.30

Laney Member

49.02 ± 0.15 49.70 ± 0.10 49.96 ± 0.08

Desertion Point Tongue of Wasatch

Cathedral Bluffs Tongue of Wasatch Formation (arkosic)

50.39 ± 0.13 50.56 ± 0.26

Wilkins Peak Member

50.70 ± 0.14 51.25 ± 0.14 New Fork Tongue of Wasatch Formation

Tipton Member Wasatch Formation

Niland Tongue of Luman Tongue Main Body of Wasatch Formation Fluvial Lacustrine (overfilled)

Evaporative (underfilled)

Deltaic Sandstone

Playa/Evaporative Lake

Freshwater Lake

Mudflat/Evaporative Lake

Alluvial plain adjacent to lake Floodplain & fluvial channels

Fluctuating Profundal (balanced) Saline to Freshwater Lake

Sandstone-mudstone unit 50.39 ± 0.13

Tuff bed with 40Ar/39Ar age in Ma

Figure 7 East-west distribution of facies in greater Green River basin. Modi¢ed slightly from Smith et al. (2003). Other sources: Sullivan (1985); Prothero (1996); Murphey et al. (1999).

Tipton deposition to fluctuating profundal conditions, whereas the Wilkins Peak records evaporative mudflat and playa conditions. The Luman and Tipton members correlate with post-Colton freshwater lacustrine margin and open lacustrine deposits of the Piceance Creek and Uinta basins; the Wilkins Peak Member correlates with nahcolite- and halite-bearing hypersaline strata in the lower part of the Parachute Creek Member of the Green River Formation in the Uinta and Piceance Creek basins ( Johnson, 1985). Single-grain 40Ar/39Ar laser-fusion ages indicate that playa and mudflat evaporative lake conditions persisted for roughly 1.5 Ma, between 51.3 and 49.7 Ma (Smith et al., 2003). An abrupt return to fluctuating profundal freshwater from saline conditions is recorded by the Laney Member and the Mahogany oil shale bed north and south of the Uinta uplift, respectively. 40 Ar/39Ar ages suggest that the abrupt shift away from evaporative conditions took place simultaneously in the greater Green River, Piceance Creek, and Uinta basins (Machlus et al., 2002; Smith et al., 2003). Individual clastic sedimentary wedges of the Wasatch Formation that impinged upon the margins of Lake Gosiute were derived from all flanks of the basin (Surdam and Stanley, 1979; Sullivan, 1985). The New Fork Tongue, which interfingers with the Tipton Member from the north and northwest, was derived from the thrustbelt reentrant at the northwest corner of the Bridger subbasin, where the Wind River uplift intersects the thrust front, and from the Wind River uplift itself (Sullivan, 1985). Lake Gosiute may have spilled to the northeast during this freshwater overfilled to balanced-fill phase of the lake (Sklenar and Andersen, 1985). The Desertion Point Tongue of the Wasatch Formation entered the basin from the southwestern reentrant of the thrust belt and interfingers northeastward with the Wilkins Peak Member, whereas the arkosic Cathedral Bluffs Tongue, which interfingers with the Wilkins Peak Member from the east, was derived from basement rocks of the Sierra Madre and Granite Mountains uplifts (Surdam and Stanley, 1980; Sullivan, 1985). The Absaroka volcanic field, which lay north of the Wind River uplift, was an important source of clastic material with a unique compositional signature for the basins lying both north and south of the Uinta uplift. The dispersal history of that clastic material demonstrates the hydrologic connection between those basins.

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108° Absaroka Volcanic Field

44°

PRB BHB 48-49 g Bi or

H n pl

U ift

W

WRB

in d

?

ive

R rU pl

Granite Mtns Uplift

BB

GDB

Lake

Sev ier

Oro gen ic

Belt

ift

IDAHO

SB

HB

Rock Gosiute Springs Uplift

48.7

Sierra Madre Uplift

WYOMING 47.6 SWB

Uinta Uplift

46.8 Lake Uinta UB

0

100 km

UTAH

Douglas Creek

42°

Arch

PB

PCB

COLORADO

Figure 8 Map of Bighorn, Greater Green River, Piceance Creek, and Uinta basins, showing uplifts that supplied sediment to the basins as discussed in the text. Progradation direction of volcanic-lithic detritus is shown by bold arrows, with ages of the arrival of the detritus from the Absaroka volcanic ¢eld in bold numbers based upon ages of tu¡ beds in Figure 7 and correlations shown in Figure 2. Bold numbers in Absaroka ¢eld indicate age of peak volcanism (Feeley and Cosca, 2003; Harlan, 2006). Solid arrows are paleocurrent vectors of Surdam and Stanley (1980), open arrows are based on general considerations of possible progradation directions (Surdam and Stanley, 1980; Johnson, 1985). Exposed basement in uplifts depicted in brown. Sedimentary basins: BB, Bridger subbasin; BHB, Bighorn basin; GDB, Great Divide subbasin; HB, Hannah basin; PB, Middle and North Park basins; PCB, Piceance Creek basin, PRB, Powder River basin; SB, Shirley basin; SWB, Sand Wash subbasin.

The Absaroka volcanic field is a large volume calc-alkaline to alkalic igneous province erupted between 55 and 45 Ma, but with peak eruptive activity from 49 to 48 Ma on the basis of high-precision 40Ar/39Ar ages (beginning near the Ypresian–Lutetion boundary; Feeley and Cosca, 2003; Harlan, 2006). A huge wedge of Middle Eocene volcanic-lithic detritus derived from the Absaroka volcanic field prograded southward across the Wind River basin (Wagon Bed Formation) and into the greater Green River basin via a gap between the Wind River and Granite Mountains uplifts (Figure 8; Surdam and Stanley, 1979, 1980). This volcaniclastic wedge (Bridger and Washakie formations) then advanced southward, crossed the east end of the Uinta uplift into the Piceance Creek basin, where it is recorded by the Uinta Formation directly overlying the Parachute Creek Member of the Green River Formation, and subsequently prograded west across the Douglas Creek arch and thence along the northern basin axis of the Uinta basin (Surdam and Stanley, 1980; Johnson, 1985). The case history of the nuclear ponded basins suggests an important interplay of climate, tectonism, and sediment supply factors on the complex mosaic of depositional environments and evolution of the basins. The Green River Formation was deposited during a pronounced warming trend of the Cenozoic extending from 59 to 52 Ma, with a peak at 52–50 Ma, termed the Early Eocene climatic optimum (EECO), followed by a return to cooler conditions that began rapidly between 50 and 48 Ma (Zachos et al., 2001). The evaporative facies of the

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Wilkins Peak and Parachute Creek members thus were deposited during the second half of the EEOC and the return to fluctuating profundal conditions took place at the end of the EEOC. It is tempting to infer that coeval drying and subsequent freshening events of Lakes Gosiute and Uinta were slightly delayed responses to global climatic trends (e.g., Roehler, 1993; Matthews and Perlmutter, 1994). Abundant evidence exists for active tectonism and uplift of basin-flanking blocks during changing conditions in Lake Gosiute. Coeval tectonism during Green River deposition is supported by: (1) asymmetric, southwardincreasing subsidence in the middle part of the Wilkins Peak Member, documented by detailed correlations along the west side of the Rock Springs uplift (Pietras et al., 2003b); (2) a nearly fourfold increase in sediment accumulation rates between the Tipton and Wilkins Peak members (Smith et al., 2003), either as a result of increased accommodation by structural elevation of the lake’s outflow sill or through increased loading by the adjacent Uinta uplift; (3) inferred coeval uplift of the north flank of the Uinta uplift during deposition of the Tipton Member (Hansen, 1984) and the south flank of the Wind River uplift during deposition of the Wilkins Peak Member (Pietras et al., 2003a); and (4) large sediment influx from the east flank and southwest corner of the basin (Sullivan, 1985). Detrital influx from the southwest probably records uplift of an integrated drainage basin that exited the thrust belt along a major cross-strike structural discontinuity (Lawton et al., 1994). Uplift of the eastern source areas may have cut off the Early Eocene outlet on the northeast (e.g., Sklenar and Andersen, 1985). The prodigious supply of easily eroded volcaniclastic debris from the Early Eocene Absaroka volcanic field overwhelmed and filled any tectonically generated accommodation space that remained in the Wind River and greater Green River basins. Because explosive volcanism strongly impacts fluvial sedimentation by providing huge volumes of mobile sediment to coeval fluvial systems (Smith, 1987), basin filling may have taken place in the absence of a climatic driving mechanism. The flood of volcanic-lithic debris containing both neovolcanic and paleovolcanic grains provides unambiguous evidence for hydrologic integration of the Wind River, greater Green River, Piceance Creek, and Uinta basins, although it does not indicate whether integration was diachronous or simultaneous. Earlier integration may be recorded by the abrupt freshening and change in water chemistry at ca. 49.5 Ma (Surdam and Stanley, 1980; Johnson, 1985; Machlus et al., 2002; Smith et al., 2003). Surdam and Stanley (1980) argued that changes in evaporite chemistry in strata below and above the Mahogany oil shale bed in the Uinta basin are consistent with expected water chemistry changes attendant upon the initial arrival of lake water from the greater Green River and Piceance Creek basins, which had different source rock distributions in their drainage basins. Because this change in evaporite composition is simultaneous in all basins within the resolution of 40Ar/39Ar laser-fusion geochronology, it may signal a net increase in runoff to the lake basins. Although runoff may have increased as a result of long-term changes in precipitation, tectonic changes in the locations of drainage divides could also have influenced hydrology of the basins (Pietras et al., 2003a). As the recipient of the collective drainage basins of Wyoming, northwestern Colorado, and probably southwestern Utah and northern Arizona, Lake Uinta likely had an outlet to the ocean in order to maintain its apparent balanced-fill or overfilled conditions (e.g., Dickinson et al., 1988). Because of the likelihood that the former thrust belt was a region of high topography in the Eocene (e.g., Chase et al., 1998; DeCelles, 2004), drainage was probably not to the west as originally suggested by Dickinson et al. (1988). No strong sediment input to the south flank of Lake Uinta is evident after the earliest Eocene ( Johnson, 1985); therefore, effluent drainage was probably southward across the San Juan basin and thence eastward near the latitude of Albuquerque and across the present Great Plains (e.g., Cather, 2004) or south along the course of the modern Rio Grande and southeast toward the Sabinas basin in northeastern Mexico (Figure 9). In summary, numerous observations indicate that tectonism, climate, and sediment supply were important factors in determining the fill of Laramide basins, in terms of both age and distribution of facies. It is clear that the relative contributions of these factors remain to be elucidated through integration of detailed regional studies and combined stratigraphic, geochronologic, and geochemical analytical approaches.

7. Tectonics of the Laramide Orogeny Spatial and temporal patterns of Laramide deformation and basin development, in combination with plate reconstructions and arc magmatic history at the continental margin, comprise the essential data for interpreting underlying mechanisms of Laramide orogenesis. The general amagmatic character of Laramide deformation, except south of the Colorado Plateau, and distribution of basins and uplifts in a cratonic setting inboard of a recently active fold and thrust belt suggest analogy with the Sierras Pampeanas province of Argentina, where the Nazca plate is being subducted subhorizontally beneath South America (Allmendinger et al., 1983; Jordan and Allmendinger, 1986). Indeed, the longest-running plate-tectonic model for Laramide deformation envisions a flat Farallon slab beneath western North America to create the far-field stresses necessary to effect the required crustal shortening (Figure 10; Dickinson and Snyder, 1978). The flat slab has been explained in terms of increased

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To Williston basin 0

300 km

44°N BasinExplanation Types ?

Perimeter basins Ponded basins

?

Axial basins InversionInversion-flank basins flank Absaroka volcanic field Fluvial drainage direction 40° N

?

To Gulf of Mexico

36° N

Albuquerque

? ? Las Cruces Tucson

32° N

116° W 112° W

108° W To Parras To Sabinas basin basin

104° W

Figure 9 Generalized drainage patterns for Laramide province. Gray arrows in southern basins represent latest Cretaceous drainage systems that £owed toward northeastern Mexico. These probably persisted until Paleogene time, but the record is not preserved. Sources: Hayes (1987); GonzaŁlez-LeoŁn and Lawton (1995); Basabilvazo (2000). Black arrows represent Middle Eocene drainages during and immediately following maximum extent of lake development of ponded basins, near end of deposition of Green River Formation. Sources: Surdam and Stanley (1980); Dickinson et al. (1988); Cather (2004).

445

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44° N 0

300 km

Approximate margins of flat slab domain

40° N

Denver

Colorado Plateau

36° N

Albuquerque

Approximate temporal direction of shortening Las Cruces

Tucson

32° N

116° W

112° W

108° W

104° W

Figure 10 Approximate distribution of £at-slab segment beneath Laramide orogen (Saleeby, 2003). Approximate average direction of temporal propagation of shortening in Laramide province is indicated by large arrow, which is parallel to inferred edges of Laramide £at-slab segment. Colors are the same as for Figure 9.

convergence rates between the Farallon and North American plates (Engebretson et al., 1984) and subduction of a buoyant aseismic ridge or oceanic plateau (Livaccari et al., 1981; Henderson et al., 1984). Analysis of paleostress directions assembled from dikes, veins, and fault slip vectors indicates that the dominant shortening direction was ENE (0681) at 75 Ma, with possible slight clockwise rotation of 15–301 at 45–40 Ma (Bird, 2002). The calculated stress directions are consistent with northeastward subduction of the Farallon plate beneath North America (Henderson et al., 1984). Early contraction that initiated basement-involved crustal deformation may have begun as early as 85 Ma through lithospheric delamination beneath southwesternmost North America as the buoyant slab segment encountered the North American plate (Saleeby, 2003). Lithospheric shearing or delamination created an end load to effect early shortening in the southwestern part of the Laramide foreland, which was then

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followed by shortening attendant upon basal shearing of the flat slab in the distal part of the Laramide province. Such a scenario is consistent with early development of the southern basins by oblique-slip along older extensional faults as the Colorado Plateau was translated northeast above the flat slab segment. Implicit in the basal traction model for crustal shortening is that a deformation front should migrate across the foreland as the flat slab segment advances beneath the continental lithosphere. Although the onset of basement uplift and basin subsidence across the Laramide foreland is not precisely constrained for reasons outlined earlier in the chapter, there is a general northwestward progression of ages apparent in the correlation chart of Figure 2. An advance of approximately 1,200 km from the McCoy basin to the Bighorn uplift between 75 and 65 Ma yields a structuralfront migration of 120 km/Ma, within the range of 90–150 km/Ma Farallon-North America convergence rates calculated for that time interval (Engebretson et al., 1984). In conclusion, flat slab model for Laramide deformation satisfies a rigorous list of tests: (1) kinematic criteria derived from an actualistic plate-tectonic setting in South America; (2) calculated plate-convergence rates between North America and the subducted Farallon plate linked to the Late Cretaceous–Paleogene shortening history in the foreland; (3) temporal and spatial distribution of Paleogene magmatism and calculated stress directions; and even (4) an oceanic plateau in the northwest Pacific basin that represents the counterpart of the subducted Laramide oceanic plateau (Henderson et al., 1984) and explains the buoyant characteristics of the subducted lithosphere. Flat subduction therefore remains a viable paradigm for the plate-tectonic origin of Laramide orogenesis.

ACKNOWLEDGMENTS I thank Andrew Barth, Ronald Blakey, William Dickinson, and Beth Welle for reviews and discussion. Although they do not agree with all of what I have written, their comments helped clarify and amplify some of my thinking. I also thank Francisco Vega, UNAM, Coyoaca´n, Me´xico D.F., and a Fulbright-Garcia Robles Grant for providing the place, time, and resources to complete this chapter.

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Bryant, B., Naeser, C. W., Marvin, R. F., and Mehnert, H. H., 1989, Upper Cretaceous and Paleogene sedimentary rocks and isotopic ages of Paleogene tuffs, Uinta basin, Utah. U.S. Geological Survey Bulletin, v. 1787J, 22 pp. Camilleri, P., Yonkee, A., Coogan, J., DeCelles, P. G., McGrew, A., and Wells, M., 1997, Hinterland to foreland transect through the Sevier orogen, northeast Nevada to north central Utah: structural style, metamorphism, and kinematic history of a large contractional orogenic wedge. Brigham Young University Geology Studies, v. 42, pp. 297–309. Carroll, A. R., and Bohacs, K. M., 1999, Stratigraphic classification of ancient lakes: balancing tectonic and climatic controls. Geology, v. 27, pp. 99–102. Cather, S. M., 2004, Laramide orogeny in central and northern New Mexico and southern Colorado, in Mack, G. H. and Giles, K. A. eds., The geology of New Mexico: a geologic history, New Mexico Geological Society (Special Publication 11), Socorro, NM, pp. 203–248. Cerveny, P. F., and Steidtmann, J. R., 1993, Fission track thermochronology of the Wind River Range, Wyoming: evidence for timing and magnitude of Laramide exhumation. Tectonics, v. 12, pp. 77–91. Chapin, C. E., and Cather, S. M., 1981, Eocene tectonics and sedimentation in the Colorado Plateau-Rocky Mountain area, in Dickinson, W. R. and Payne, W. D. eds., Relations of tectonics to ore deposits in the southern Cordillera, Tucson, AZ, pp. 173–198. Chase, C. G., Gregory, K. M., Parrish, J. T., and DeCelles, P. G., 1998, Topographic history of the western Cordillera of North America and the etiology of climate, in Crowley, T. J. and Burke, K. eds., Tectonic boundary conditions for climate reconstructions, Oxford Monographs on Geology and Geophysics, Oxford, v. 39, pp. 73–99. Christensen, A. 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Goldstrand, P. M., 1992, Evolution of Late Cretaceous and Early Tertiary basins of southwest Utah based on clastic petrology. Journal of Sedimentary Petrology, v. 62, pp. 495–507. Goldstrand, P. M., 1994, Tectonic development of Upper Cretaceous to Eocene strata of southwestern Utah. Geological Society of America Bulletin, v. 106, pp. 145–154. Gonza´lez-Leo´n, C. M., and Lawton, T. F., 1995, Stratigraphy, depositional environments, and origin of the Cabullona basin, northeastern Sonora, in Jacques-Ayala, C., Gonza´lez-Leo´n, C. M., and Rolda´n-Quintana, J. eds., Studies on the Mesozoic of Sonora and adjacent areas, Geological Society of America (Special Paper 301), Boulder, CO, pp. 121–142. Gradstein, F. M., Ogg, J. G., Smith, A. G., Bleeker, W., and Lourens, L. J., 2004, A new Geologic Time Scale with special reference to Precambrian and Neogene. Episodes, v. 27, pp. 83–100. Gries, R., 1983a, Oil and gas prospecting beneath Precambrian of foreland thrust plates in Rocky Mountains. American Association of Petroleum Geologists Bulletin, v. 67, pp. 1–28. Gries, R., 1983b, North-south compression of Rocky Mountain foreland structures, in Lowell, J. D. and Gries, R. eds., Rocky Mountain foreland basins and uplifts, Rocky Mountain Association of Geologists, Denver, CO, pp. 9–32. Guzman, E. J., and de Cserna, Z., 1963, Tectonic history of Mexico, in Childs, O.E., and Beebe, B.W. eds., Backbone of the Americas— tectonic history from pole to pole, American Association of Petroleum Geologists, Tulsa, OK, Memoir 2, pp. 113–129. Hagen, E. S., Schuster, M. W., and Furlong, K. P., 1985, Tectonic loading and subsidence of intermontane basins: Wyoming foreland province. Geology, v. 13, pp. 529–592. Hansen, W. R., 1984, Post-Laramide tectonic history of the eastern Uinta Mountains, Utah, Colorado, and Wyoming. The Mountain Geologist, v. 21, pp. 5–29. Harlan, S. 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A., 2000, Structural geology and Laramide tectonics of the Little Hatchet Mountains, southwestern New Mexico, New Mexico Geological Society, Socorro, NM, Guidebook 51, pp. 109–116. Inman, K. F., 1987, Depositional environments and sandstone petrography of Cretaceous sedimentary rocks, Adobe Canyon, Santa Rita Mountains, southeastern Arizona, in Dickinson, W. R. and Klute, M. A. eds., Mesozoic rocks of southern Arizona and adjacent areas, Arizona Geological Society, Tucson, AZ, Digest 18, pp. 301–314. Iriondo, A., Premo, W. R., Martı´nez-Torres, L. M., Budahn, J. R., Atkinson, W. W., Jr., Siems, D. F., and Guara´s-Gonza´les, B., 2004, Isotopic, geochemical, and temporal characterization of Proterozoic basement rocks in the Quitovac region, northwestern Sonora, Mexico: implications for the reconstruction of the southwestern margin of Laurentia. Geological Society of America Bulletin, v. 116, pp. 154–170. Isby, J. S., and Picard, M. D., 1985, Depositional setting of Upper Cretaceous–lower Tertiary Currant Creek Formation, north-central Utah, Utah Geological Association Salt Lake city, UT, Publication 12, pp. 39–49. Jacques-Ayala, C., 1999, Stratigraphy of El Chanate Group (Late Cretaceous) and its implications for the tectonic evolution of northwestern Sonora, Mexico. Revista Mexicana de Ciencias Geolo´gicas, v. 16, pp. 97–120. Jacques-Ayala, C., Barth, A. P., Wooden, J. L., and Jacobson, C. E., 2005, Provenance and U-Pb geochronology of the Late Cretaceous El Chanate Group, NW Sonora, Mexico, GEOS, Union Geofisica Mexicana, A.C., Boletin Informativo, Epoca II, v. 25, pp. 160–161. Johnson, R. C., 1985, Early Cenozoic history of the Uinta and Piceance Creek basins, Utah and Colorado, with special reference to the development of Eocene Lake Uinta, in Flores, R. M. and Kaplan, S. S. eds., Cenozoic paleogeography of the west-central United States, Rocky Mountain Section, Society of Economic Paleontologists and Mineralogists (SEPM), Denver, CO, pp. 247–276. Jordan, T. E., and Allmendinger, R. W., 1986, The Sierras Pampeanas of Argentina: a modern analogue of Rocky Mountain foreland deformation. American Journal of Science, v. 286, pp. 737–764. Karlstrom, K. E., Amato, J. M., Williams, M. L., Heizler, M., Shaw, C., Read, A., and Bauer, P., 2004, Proterozoic evolution of the New Mexico region: a synthesis, in Mack, G. H. and Giles, K. A. eds., The geology of New Mexico: a geologic history, New Mexico Geological Society (Special Publication 11), Socorro, NM, pp. 1–34. Keefer, W. R., 1965, Stratigraphy and geologic history of the uppermost Cretaceous, Paleocene, and lower Eocene rocks in the Wind River basin, Wyoming, U.S. Geological Survey, Denver, CO, Professional Paper 495A, 77 pp. Kelley, S. A., and Chapin, C. E., 1995, Apatite fission-track thermochronology of southern Rocky Mountain-Rio Grande rift-western High Plains provinces, New Mexico Geological Society, Socorro, NM, 46th Field Conference Guidebook, pp. 87–96. Kraus, M. J., 1987, Integration of channel and floodplain suites, II. Vertical relations of alluvial paleosols. Journal of Sedimentary Petrology, v. 57, pp. 602–612. Lawton, T. F., 1983, Late Cretaceous fluvial systems and the age of foreland uplifts in central Utah, in Lowell, J. D. and Gries, R. eds., Rocky Mountain foreland basins and uplifts, Rocky Mountain Association of Geologists, Denver, CO, pp. 181–199. Lawton, T. F., 1986, Fluvial systems of the Upper Cretaceous Mesaverde Group and Paleocene North Horn Formation, central Utah: a record of transition from thin-skinned to thick-skinned deformation in the foreland region, in Peterson, J. A. ed., Paleotectonics and sedimentation in the Rocky Mountain region, United States, American Association of Petroleum Geologists, Tulsa, OK, Memoir 41, pp. 423–442. Lawton, T. F., 1994, Tectonic setting of Mesozoic sedimentary basins, Rocky Mountain region, United States, in Caputo, M. V., Peterson, J. A., and Franczyk, K. A. eds., Mesozoic systems of the Rocky Mountain region, USA, Society for Sedimentary Geology (SEPM), Denver, CO, pp. 1–25. Lawton, T. F., 2000, Inversion of Late Jurassic–Early Cretaceous extensional faults of the Bisbee basin, southeastern Arizona and southwestern New Mexico, New Mexico Geological Society, Socorro, NM, Guidebook 51, pp. 95–102.

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Lawton, T. F., and Trexler, J. H., Jr. 1991, Piggyback basin in the Sevier orogenic belt, Utah: implications for development of the thrust wedge. Geology, v. 19, pp. 827–830. Lawton, T. F., Basabilvazo, G. T., Hodgson, S. A., Wilson, D. A., Mack, G. H., McIntosh, W. C., Lucas, S. G., and Kietzke, K. K., 1993a, Laramide stratigraphy of the Little Hatchet Mountains, southwestern New Mexico. New Mexico Geology, v. 15, pp. 9–18. Lawton, T. F., Talling, P. J., Hobbs, R. S., Trexler, J. H., Jr., Weiss, M. P., and Burbank, D. W., 1993b, Structure and stratigraphy of Upper Cretaceous and Paleogene strata (North Horn Formation), eastern San Pitch Mountains, Utah–Sedimentation at the front of the Sevier orogenic belt, U.S. Geological Survey, Denver, CO, Professional Paper 1787-II, 33 pp. Lawton, T. F., Boyer, S. E., and Schmitt, J. G., 1994, Influence of inherited taper on structural variability and conglomerate distribution, Cordilleran fold and thrust belt, western United States. Geology, v. 22, pp. 339–342. Lawton, T. F., Pollock, S. L., and Robinson, R. A. J., 2003, Integrating sandstone petrology and nonmarine sequence stratigraphy: application to the Late Cretaceous fluvial systems of southwestern Utah, U.S.A. Journal of Sedimentary Research, v. 73, pp. 389–406. Lawton, T. F., Gonza´lez-Leo´n, C. M., Amato, J. M., Gehrels, G. E., Heizler, M., Iriondo, A., Leggett, W. J., Mauel, D. J., and Peryam, T. C., 2008, Contrasting tectonic regimes recorded by sedimentary basins associated with Jurassic and Cretaceous are systems, Sonora, Mexico: l er congreso sabre la Evolucio´n Geolo´gica y Ecolo´gica del Noroeste de Me´xico, Libro de Resu´menes, Hermosillo, Sonora, Mexico, pp. 105–107. Livaccari, R. F., Burke, K., and Sengo¨r, A. M. C., 1981, Was the Laramide orogeny related to subduction of an oceanic plateau? Nature, v. 289, pp. 276–278. Lucas, S. G., Basabilvazo, G. T., and Lawton, T. F., 1990, Late Cretaceous dinosaurs from the Ringbone Formation, southwestern New Mexico. Cretaceous Research, v. 11, pp. 1–7. Machlus, M., Hemming, S. R., Olsen, P. E., and Christie-Blick, N., 2002, 40Ar/39Ar ages of selected tuffs of the Green River Formation; implications for interbasin and intrabasin correlations. Geological Society of America Abstracts with Programs, v. 34(6), pp. 558. Mann, J. D., 1995. Stratigraphy, petrology, and depositional environments of the Upper Cretaceous Rucker Formation, southeastern Arizona, M.S. thesis, New Mexico State University, Las Cruces, 114 pp. Marshak, S., Karlstrom, K., and Timmons, J. M., 2000, Inversion of Proterozoic extensional faults: an explanation for the patterns of Laramide and Ancestral Rockies intracratonic deformation. Geology, v. 28, pp. 735–738. Matthews, M. D., and Perlmutter, M. A., 1994, Global cyclostratigraphy: an application to the Eocene Green River basin, in de-Boer, P. L. and Smith, D. G. eds., Orbital forcing and cyclic sequences, International Association of Sedimentologists (Special Publication 19), Oxford, pp. 459–481. McDowell, F. W., Roldan-Quintana, J., and Connelly, J. N., 2001, Duration of Late Cretaceous–Early Tertiary magmatism in east-central Sonora, Mexico. Geological Society of America Bulletin, v. 113, pp. 521–531. McQuarrie, N., and Chase, C. G., 2000, Raising the Colorado Plateau. Geology, v. 28, pp. 91–94. Murphey, P. C., Lester, A., Bohor, B., Robinson, P., Evanoff, E., and Larson, E., 1999, 40Ar/39Ar dating of volcanic ash deposits in the Bridger Formation (Middle Eocene), southwestern Wyoming. Geological Society of America Abstracts with Programs, v. 31(7), 233 pp. Nummedal, D., 2004, Tectonic and eustatic controls on Upper Cretaceous stratigraphy of northern New Mexico, in Mack, G. H. and Giles, K. A. eds., The Geology of New Mexico: a geologic history, New Mexico Geological Society (Special Publication 11), Socorro, NM, pp. 169–182. Perry, W. J., Jr., Haley, J. C., Nichols, D. J., Hammons, P. M., and Ponton, J. D., 1988, Interactions of Rocky Mountain foreland and Cordilleran thrust belt in Lima region, southwest Montana, in Schmidt, C. J. and Perry, W. J., Jr. eds., Interaction of the Rocky Mountain foreland and Cordilleran thrust belt, Geological Society of America, Boulder, CO, Memoir 171, pp. 267–290. Pietras, J. T., Carroll, A. R., and Rhodes, M. K., 2003a, Lake basin response to tectonic drainage diversion: Eocene Green River Formation, Wyoming. Journal of Paleolimnology, v. 30, pp. 115–125. Pietras, J. T., Carroll, A. R., Singer, B. S., and Smith, M. E., 2003b, 10 k.y. depositional cyclicity in the Early Eocene: stratigraphic and 40 Ar/39Ar evidence from the lacustrine Green River Formation. Geology, v. 31, pp. 593–596. Prothero, D. R., 1996, Summary, in Prothero, D. R. and Emry, R. J. eds., The terrestrial Eocene–Oligocene transition in North America, Cambridge University Press, Cambridge, pp. 664–683. Roberts, E. 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F., 1997, Structural kinematics and depositional history of a Laramide uplift-basin pair in southern New Mexico: implications for development of intraforeland basins. Geological Society of America Bulletin, v. 109, pp. 1389–1401.

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Sklenar, S. E., and Andersen, D. W., 1985, Origin and early evolution of an Eocene lake system with the Washakie basin of southwestern Wyoming, in Flores, R. M. and Kaplan, S. S. eds., Cenozoic paleogeography of the west-central United States, Rocky Mountain Section, Society of Economic Paleontologists and Mineralogists (SEPM), Denver, CO, pp. 231–245. Smith, G. A., 1987, The influence of explosive volcanism on fluvial sedimentation: the Deschutes Formation (Neogene) in central Oregon. Journal of Sedimentary Petrology, v. 57, pp. 613–629. Smith, M. E., Singer, B., and Carroll, A., 2003, 40Ar/39Ar geochronology of the Eocene Green River Formation, Wyoming. Geological Society of America Bulletin, v. 115, pp. 549–565. Smithson, S. B., Brewer, J., Kaufman, S., Oliver, J. E., and Hurich, C., 1978, Nature of the Wind River Thrust, Wyoming, from COCORP deep-reflection data and from gravity data. Geology, v. 6, pp. 648–652. Smoot, J. P., 1983, Depositional subenvironments in an arid closed basin: the Wilkins Peak Member of the Green River Formation (Eocene), Wyoming, U.S.A. Sedimentology, pp. 801–827. Stanley, K. O., and Collinson, J. W., 1979, Depositional history of Paleocene–Lower Eocene Flagstaff Limestone and coeval rocks, central Utah. American Association of Petroleum Geologists Bulletin, v. 63, pp. 311–323. Steidtmann, J. R., 1971, Origin of the Pass Peak Formation and equivalent Early Eocene strata, central western Wyoming. Geological Society of America Bulletin, v. 82, pp. 159–176. Sullivan, R., 1985, Origin of lacustrine rocks of Wilkins Peak Member, Wyoming. American Association of Petroleum Geologists Bulletin, v. 69, pp. 913–922. Surdam, R. C., and Stanley, K. O., 1979, Lacustrine sedimentation during the culminating phase of Eocene Lake Gosiute, Wyoming (Green River Formation). Geological Society of America Bulletin, v. 90, pp. 93–110. Surdam, R. C., and Stanley, K. O., 1980, Effects of changes in drainage-basin boundaries on sedimentation in Eocene Lakes Gosiute and Uinta of Wyoming, Utah, and Colorado. Geology, v. 8, pp. 135–139. Thorman, C. T., and Drewes, H., 1980, Geologic map of the Victorio Mountains, Luna County, southwestern New Mexico, U. S. Geological Survey, Map MF-1175, scale 1:24,000. Tosdal, R. M., 1990, Constraints on the tectonics of the Mule Mountains thrust system, southeast California and southwest Arizona. Journal of Geophysical Research, v. 95, pp. 20,025–20,048. Tosdal, R. M., and Stone, P., 1994, Stratigraphic relations and U-Pb geochronology of the Upper Cretaceous upper McCoy Mountains Formation, southwestern Arizona. Geological Society of America Bulletin, v. 106, pp. 476–491. Weiss, M. P., and Warner, K. N., 2001, The Crazy Hollow Formation (Eocene) of central Utah. Brigham Young University Geology Studies, v. 46, pp. 143–161. Woodward, L. A., Kaufman, W. H., and Anderson, J. B., 1972, Nacimiento fault and related structures, northern New Mexico. Geological Society of America Bulletin, v. 83, pp. 2383–2396. Yin, A., and Ingersoll, R. V., 1997, A model for evolution of Laramide axial basins in the southern Rocky Mountains, U.S.A. International Geology Review, v. 39, pp. 1113–1123. Young, J. R., McMillan, N. J., Lawton, T. F., and Esser, R. P., 2000, Volcanology, geochemistry, and structural geology of the Upper Cretaceous Hidalgo Formation, southwestern New Mexico, New Mexico Geological Society, Socorro, NM, Guidebook 51, pp. 149–157. Zachos, J., Pagani, M., Sloan, L., Thomas, E., and Billups, K., 2001, Trends, rhythms, and aberrations in global climate 65 Ma to present. Science, v. 292, pp. 685–693.

CHAPTER 13

Sverdrup Basin Ashton Embry and Benoit Beauchamp

Contents 1. Introduction 2. Geological Setting 3. Depositional and Tectonic History 3.1. Phase 1: From mountains to depressions (Early Carboniferous–Early Late Carboniferous) 3.2. Phase 2: Repeated quiescence and inversion (Late Carboniferous–Early Permian) 3.3. Phase 3: Passive and cold (Middle–Late Permian) 3.4. Phase 4: Filling the deep basin (Triassic) 3.5. Phase 5: Shallow seas (latest Triassic–earliest Cretaceous) 3.6. Phase 6: Rejuvenation (Early Cretaceous–earliest Late Cretaceous) 3.7. Phase 7: Quiescence (Late Cretaceous) 3.8. Phase 8: Fragmentation and uplift (Paleocene–Eocene) 4. Tectonic Episodes 5. Economic Geology 5.1. Petroleum 5.2. Coal 6. Summary References

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Abstract The Sverdrup Basin is located in the Canadian Arctic Islands. It is 1,000 km by 350 km and is ﬁlled with up to 13 km of Carboniferous to Paleogene strata. The basin initially developed in Early Carboniferous as a rift basin upon highly deformed Early Paleozoic strata of the Ellesmerian Orogenic Belt. The development of the basin can be broken into eight phases, each being characterized by a distinctive combination of tectonic, depositional and climatic regimes and separated by episodes of widespread uplift and basin reorganization. The Upper Paleozoic strata are up to 5 km thick and are characterized by a distinct shelf to deep basin topography. Carbonate strata dominated the shelf until Middle Permian and were supplanted by siliciclastics and chert in Middle and Late Permian when the climate cooled. Triassic siliciclastics are up to 4 km thick and they ﬁlled the deep, central basin by Late Triassic. From latest Triassic to earliest Cretaceous the basin was occupied by shallow siliciclastics shelves and up to 2 km of strata accumulated. Renewed rifting in Early Cretaceous resulted in a thick succession (2 km) of Early Cretaceous nonmarine to shallow marine strata with units of basalts in the northeast. Widespread diabase sill and dyke intrusion, likely related to the Alpha Ridge Plume and the opening of the Amerasia Ocean Basin, occurred at this time. Following an interval of low subsidence and low sediment supply in the Late Cretaceous, the basin began to be deformed in earliest Paleocene by the Eurekan Orogeny driven by the counterclockwise rotation of Greenland. Local foreland basins developed and contain up to 3 km of Paleocene–Eocene strata. In Late Eocene the basin was uplifted and deformed by faulting and folding with deformation decreasing southwestwards. Eighteen oil and gas ﬁelds have been discovered in Eurekan anticlines and potential prospects include traps associated with Eurekan structures, salt domes, reefs and prominent unconformities. Widespread petroleum source rocks are documented in Middle and Upper Triassic strata and likely occur with other stratigraphic intervals from Carboniferous to Lower Cretaceous.

Sedimentary Basins of the World, Volume 5 ISSN 1874-5997, DOI 10.1016/S1874-5997(08)00013-0

r 2008 Elsevier B.V. All rights reserved.

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1. Introduction The Sverdrup Basin, a major Carboniferous to Paleogene depocentre in the Canadian Arctic Islands, was recognized during Operation Franklin (1955), when the geology of the Arctic Islands was ﬁrst systematically studied by the Geological Survey of Canada (Fortier et al., 1963). The basin occupies much of the Queen Elizabeth Islands and stretches for about 1,300 km between northern Ellesmere Island on the northeast and Prince Patrick Island on the southwest. It is up to 350 km across (Figure 1). About half of the 300,000 km2 area of the basin is land with the remainder being inter-island channels. The basin ﬁll consists of Carboniferous to Eocene strata that are estimated to be up to 13 km thick (Balkwill, 1978). The eastern portion of the basin was uplifted and deformed in the Late Paleogene and is now a mountainous area. To the west, deformation was much less and the topography is subdued (Figure 2). The folding and thrusting of the eastern portion have deformed the original shape of the basin and palinspastic restorations have not been attempted. Ofﬁcers of the Geological Survey have described the outcropping strata in a series of publications; the pre1988 results are summarized in the Arctic Islands DNAG volume (Trettin, 1991). Publications such as Harrison (1995), Beauchamp (1995) and Beauchamp et al. (2001) provide recent information on the Late Paleozoic succession, building from the pioneering work of Thorsteinsson (1974), Mayr (1992) and Davies and Nassichuk (1991a). Embry (1991a) summarized the stratigraphy and depositional history of the Mesozoic succession. Ricketts (1994) and Ricketts and Stephenson (1994) are the most comprehensive summaries of Tertiary stratigraphy and paleogeographic evolution. Oil exploration in the basin began in 1968, and between 1969 and 1986 120 wells were drilled into Sverdrup Basin strata. Tens of thousands of kilometers of seismic lines were shot in the basin although most of these lines have not been studied in detail. The exploratory activity resulted in 18 discoveries with natural gas being the primary hydrocarbon found.

Figure 1 Outline of Sverdrup Basin with basin axis. The black dots represent the 120 wells drilled in the basin. The main source areas lay to the south and east as indicated by arrows. A low-lying land area to the north (Crockerland) was a relatively minor source area. TH,Tanquary High.

Sverdrup Basin

Figure 2

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Deformation zones of the basin with intensity of deformation decreasing westward.

The availability of a large amount of surface and subsurface data has allowed a reasonably good understanding of the regional depositional and tectonic history of the basin. As described below, the development of the basin has been subdivided into eight phases, each of which is characterized by a unique combination of tectonic, depositional and climatic factors.

2. Geological Setting The Phanerozoic geology of the Canadian Arctic Islands was summarized by Trettin (1989). From Cambrian through Early Devonian a passive margin basin, the Franklinian Basin occupied the area. It was dominated by thick, shelf carbonates and argillaceous basin deposits. The Ellesmerian Orogen progressively deformed the Franklinian Basin from Late Silurian to the end of the Devonian and a huge siliciclastic wedge up to 10 km thick was deposited in front of the southwesterly advancing orogen (Embry, 1991b). The Sverdrup Basin originated in the Early Carboniferous and it developed as a rift basin on the SiluroDevonian Ellesmerian Orogenic Belt (Balkwill, 1978). The Ellesmerian Orogenic Belt of the Arctic Islands was part of an extensive Silurian–Devonian orogenic system that extended thousands of kilometers from the Acadian Orogen of the Appalachians in the south, through the Caledonides of northern Europe and Greenland to the southwestern end of the Ellesmerian belt in the northern Yukon. Notably, Carboniferous rift basins subsequently developed along much of this composite Siluro-Devonian orogenic zone (Ziegler, 1988). Such a regional shift from compression to extension suggests that a global plate-tectonic reorganization occurred in Early Carboniferous and the occurrence of such a large-scale tectonic reorganization over much of North America has been noted in other chapters of this book (e.g., see Miall and Blakey, Chapter 1). The Sverdrup Basin is the largest of these extensional ‘‘successor’’ basins that formed along the Caledonian/Ellesmerian orogenic belt during this major global restructuring. The development of the basin can be conveniently described in terms of eight phases with each phase having a distinctive combination of tectonic, depositional and climatic inﬂuences (Figure 3). The strata of each phase are bound by major, angular unconformities and correlative surfaces, indicating that a signiﬁcant and short-lived tectonic episode initiated and terminated each phase. Lesser magnitude, tectonic episodes, marked by unconformities on the basin margins, occur within these phases and notable changes in depositional regimen often occur across such discontinuities. The stratigraphy of the basin is summarized in three stratigraphic charts and a regional cross-section for the Mesozoic strata (Figures 4, 5, 6 and 7). Most of these major, phase-bounding unconformities correlate plus or minus a stage with the widespread unconformities which Sloss (1988) used to subdivide the Phanerozoic succession of the cratonic North America into 6 sequences with 16 component subsequences (see Burgess, Chapter 2). The subsequences of Sloss, which approximate the eight phases discussed in this report, are � Phase 1. Kaskaskia II � Phase 2. Absaroka I

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Figure 3

The eight phases of development with the characteristic features and timing of each phase.

� � � � � �

3. 4. 5. 6. 7. 8.

Phase Phase Phase Phase Phase Phase

Absaroka II lower portion Absaroka III upper portion Absaroka III and Zuni I Zuni II Zuni III Tejas 1.

The main sediment source areas for the basin were the adjacent Ellesmerian Fold Belt immediately to the east and south of the basin, as well as the more distant Greenland and Canadian cratonic areas. The main bedrock in all these source areas consisted of Devonian siliciclastic strata previously derived from the Caledonian/Ellesmerian Orogenic Belt (Patchett et al., 2004). A small land area, named Crockerland, lay to the northwest of the basin and it provided smaller amounts of siliciclastic sediments (Embry, 1992) (Figure 1). The evidence for the existence of the land area consists mainly of the occurrence of shallow-water, sandstonedominant strata of Permian to Jurassic age along the northwest margin of the basin. Such shallow-water lithologies change facies to deeper water argillaceous strata southwards. The volumes of such northerly derived strata leave little doubt as to the existence of a substantial land area to the north of the Arctic Islands from Carboniferous through Early Jurassic. Field work on northern Axel Heiberg Island in the late 1990s conﬁrmed the presence of numerous northerly derived, sandstone units ranging in age from Permian to Middle Jurassic (Embry, 1997). A prominent tectonic arch, the Tanquary High, extends into the northeastern portion of the basin (Nassichuk and Christie, 1969; Maurel, 1989) (Figure 1). It was a signiﬁcant positive feature until latest Triassic, and Carboniferous to Norian strata are truncated on the ﬂanks of the arch. Latest Triassic (Rhaetian) strata overlie the crest of the arch and it was not a positive tectonic feature from Rhaetian onwards. The adjacent Amerasia Basin is currently interpreted to have undergone a rift phase from Early Jurassic to earliest Cretaceous followed by a period of seaﬂoor spreading in Early Cretaceous (Embry and Dixon, 1994; Mickey et al., 2002). As will be discussed, the tectonic development of this nearby oceanic basin had a signiﬁcant inﬂuence on the tectonic and paleogeographic development of the Sverdrup Basin.

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Figure 4 Stratigraphic chart showing Carboniferous and Permian units of Sverdrup Basin. Note sequence-bounding unconformities at basin margin continuous sedimentary record in distal areas. Stratigraphic position and direction of tectonic pulses also indicated. Modi¢ed from Beauchamp et al. (2001).

3. Depositional and Tectonic History 3.1. Phase 1: From mountains to depressions (Early Carboniferous–Early Late Carboniferous) Following the widespread erosion of the Ellesmerian Orogen, rifting in a general north-south direction led to the creation of the Sverdrup Basin in part through extensional reactivation of older thrust faults in the Franklinian Basement (Harrison, 1995) (Figure 8). The earliest rift development is recorded in the Emma Fiord Formation, a unit known from a handful of outcrops on Devon, Axel Heiberg and Ellesmere islands, ranging in thickness from a few tens of meters to more than 700 m (Davies and Nassichuk, 1988). Vise´an (Early Carboniferous) Emma Fiord sedimentation occurred in a range of lacustrine (shales, carbonates), ﬂuvial (sandstones, conglomerates) and marginally marine (carbonates) environment. Lower Carboniferous (Tournaisian) strata in potentially larger subbasins may be concealed in the subsurface. The Emma Fiord Formation is characteristically black due to its high organic content. It was deposited in a warm and humid and paleoclimate as shown by a rich palyoﬂora (Davies and Nassichuk, 1988). It is locally an oil shale with signiﬁcant source rock potential (see below). The second phase of rift development, evidenced by mappable facies distribution of syn-rift conglomerates (The´riault et al., 1993) and observation of growth faults on seismic proﬁles (Harrison, 1995; Beauchamp et al.,

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Figure 5

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Triassic--Jurassic stratigraphic chart.

2001) (Figure 8), led to a signiﬁcant enlargement of the rift system. This is recorded in the red-colored Borup Fiord Formation (Thorsteinsson, 1974), which lies unconformably on Emma Fiord strata, or directly on the Franklinian basement with a profound angular unconformity (Figure 4). A variety of rift-related ﬂuvial environments and their deposits are recognized including alluvial fans and braid-plain mass-ﬂow and streamchannel conglomerates with large Franklinian clasts derived from subbasin-bounding highs, to sheetﬂood sandstones and mudstones deposited along the axis of tectonic depressions (The´riault et al., 1993). Minor evaporites locally occur, which, in addition to the red coloration and the widespread development of caliches, indicate a shift to a semi-arid climate. Volcanics (Audhild volcanics) locally occur within or immediately above the Borup Fiord Formation. The Borup Fiord succession passes upward into a marine limestone of Serpukhovian age (late Early Carboniferous). Shortly thereafter a signiﬁcant base-level drop, coinciding with the Mississippian– Pennsylvanian boundary (Early–Late Carboniferous boundary), led to widespread erosion of Borup Fiord strata and to the progradation of a wedge of marine sandstone in the distal basin axial areas where it lies in the lowest part of the Otto Fiord Formation (Mayr, 1992). The third phase of rifting, also evidenced by growth faults on seismic proﬁles (Figure 8) (Harrison, 1995; Beauchamp et al., 2001) and the distribution of syn-rift conglomerates and evaporites, occurred during the Bashkirian–Early Moscovian interval (The´riault, 1991) and led to both a signiﬁcant deepening and to a major enlargement of the Sverdrup Basin tens of kilometers outboard of the initial rift conﬁguration. Fault-controlled subsidence in the axial area led to the widespread invasion of marine waters resulting in a cyclic succession of

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Figure 6

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Cretaceous--Paleogene stratigraphic chart.

open to restricted marine carbonates and subaqueous evaporites (gypsum, anhydrite, halite) (Figure 4). These rocks belong to the Otto Fiord Formation, which is up to 400 m thick in the outcrop belt, but a greater thickness was probably deposited along the basin axis, as suggested by the occurrence of hundreds of diapiric structures piercing the overlying succession to the west (Nassichuk and Davies, 1980). Contemporaneous sedimentation at the basin margin was similar to that of the earlier Borup Fiord Formation: red-colored conglomerates and sandstones of alluvial fan to braided-river origin deposited in a series of fault-bounded tectonic depressions and passing upward into marginal marine carbonates and rare evaporites. Together, the Otto Fiord Formation and the correlative lower Canyon Fiord Formation constitute the transgressive systems track of a broad unconformitybounded sequence, the maximum ﬂooding surface of which marks the end of orthogonal rifting in the Sverdrup Basin (i.e., extension perpendicular to major basement structures) (Beauchamp et al., 2001).

3.2. Phase 2: Repeated quiescence and inversion (Late Carboniferous–Early Permian) The second phase of basin development was marked by four episodes of tectonic quiescence characterized by passive and uniform subsidence separated by shorter intervals of regional uplift and fault-controlled differential subsidence, during which some of the previously developed half-grabens continued to grow, while others ceased to be active, and new depressions, fault-bounded highs and broad ﬂexures were developed. This suggests a reorganization of tectonic stresses, both in terms of magnitude and orientation, as orthogonal north-south rifting of the previous phase gave way to a more complex history of transtension and transpression along a northwestsoutheast principal compressional stress direction, and a secondary northeast-southwest extensional direction (Figure 4), thus oblique to most of the previous rift structures (Beauchamp et al., 2001). The onset of this second phase was marked by a rapid base-level rise. This event is recorded by the drowning of the entire Sverdrup Basin area, a complete reconnection with the open ocean, a major invasion of the sea tens of kilometers inland (middle limestone member, Canyon Fiord Formation), and the offshore growth of huge keep-up carbonate reef-mounds (Tellevak member of Hvitland Peninsula Formation), some of which attained

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Regional stratigraphic cross-section of Mesozoic strata (from Embry, 1991a).

Figure 8 Subsurface north-south cross-section on Prince Patrick Island showing stratigraphic sequences. Note rift-related growth faults in early part of succession. Based on interpretation of seismic pro¢le tying Jameson Bay C-31 and Satellite F-68 wells. Modi¢ed from Beauchamp et al. (2001).

thicknesses of up to 600 m before being drowned themselves (Davies and Nassichuk, 1991b). This gain in accommodation likely resulted from crustal collapse associated with the cessation of rifting and fault-controlled subsidence. The next 12 million years (Late Moscovian to Kasimovian) were marked by slower, regional, uniform and passive subsidence which allowed a carbonate platform (Nansen Formation) to prograde basinward for signiﬁcant distances. Nansen shelf carbonates are typically arranged in a series of high-frequency sequences, or cyclothems, which may have been controlled by glacio-eustatic ﬂuctuations due to the advance and retreat of contemporaneous Gondawana glaciers, as suggested by the occurrence of similar cycles around the world (Heckel, 1986). Nansen carbonates pass landward into basin-fringing sandstones deposited in shallow subtidal to coastal plain environments (upper clastic member, Canyon Fiord Formation) and basinward into large shelf-edge reefs that are up to 1.5 km thick. Correlative, yet substantially thinner, slope to basinal turbidites and mudrocks are contained in the Hare Fiord Formation (Figure 9). These observations, plus the occurrences of a wide variety of tropical-like biotic and abiotic carbonate components attest for a warm marine environment (Beauchamp and Desrochers, 1997).

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Figure 9 Late Carboniferous paleogeography with shallow water carbonates with reefs bordering a deep central basin.

A signiﬁcant base-level drop prior to the end carboniferous (Kasimovian–Gzhelian boundary) led to a widespread subaerial unconformity and coincides with the renewal of fault-controlled differential subsidence. While subsidence remained generally rapid regionally, and led to an overall deepening-upward succession, local highs developed leading to the creation of a number of subbasins. One such subbasin in west-central Ellesmere Island became isolated during the Early Permian and was the site of increasingly deep-water subaqueous evaporite sedimentation (upper Antoinette and Mount Bayley formations). This time marks also the onset of uplift along the Tanquary High, a major positive element that remained episodically active until the end of the Triassic. Faultcontrolled subsidence ceased abruptly during the Late Asselian, once again associated with rapid base-level rise, keep-up reef-mound development (Tolkien reefs of Beauchamp and Olchowy, 2003), cessation of evaporite sedimentation, creation of a major maximum ﬂooding surface and outpouring of volcanism on Axel Heiberg Island (upper Nansen Formation). Carbonate progradation resumed as recorded in up to 500 m of highly fossiliferous strata in the upper Nansen and Tanquary formations. The next major base-level drop mid-way through the Sakmarian not only marked the renewal of differential, fault-controlled subsidence and uplift, it also coincided with a rapid climatic change as warm water carbonates of the Nansen Formation gave way to cool- to cold-water carbonates of the overlying Raanes Formation and younger formations (Beauchamp and Henderson, 1994), an abrupt climate change event that has left a deep imprint around the globe (Beauchamp and Baud, 2002). The Raanes Formation (100–250 m) marks the onset of cool-water sedimentation in the Sverdrup Basin as shown by the extensive deposition of carbonates with impoverished biota (heterozoan biota of James, 1997). Two progradational packages, the up to 300-m thick Artinskian Great Bear Cape Formation (and correlative deep-water Trappers Cove Formation), and overlying, yet unnamed, Kungurian carbonates (up to 250 m thick) and deep-water correlatives also display cool water, heterozoan carbonates. Kungurian carbonates pass landward into a thick succession of shallow subtidal sandstones and coastal plain sediments, interspersed with black coaly seams (Sabine Bay Formation), thus indicating a short-lived shift to more humid conditions. Both the Artinskian and Kungurian successions prograded basinward during episodes of tectonic quiescence and passive subsidence following short-lived intervals of faulting and folding, volcanism (Esayoo Formation), rapid subsidence associated with tectonic stress relaxation and creation of maximum ﬂooding surfaces. Phase 2 ended in a series of events that changed forever the proﬁle of the Sverdrup Basin succession. First a signiﬁcant base-level drop led to the widespread erosion of Sabine Bay sandstones and correlative carbonates and to extensive basinward progradation of both lithologies. Signiﬁcant faulting, mostly compressional in nature, followed and this led to more erosion of Kungurian strata at the basin margin, as well as local folding and erosional stripping of even older formations. These events, of variable magnitude and extent in different parts of the basin, are collectively termed the ‘‘Melvillian Disturbance’’ (Thorsteinsson, 1974), which coincides with the Early– Middle Permian boundary (Figure 4). Up until the Early Cretaceous, tectonic events, although undeniably present, left a far subtler imprint, mostly in the form of broad regional uplifts, than during the previous 80 million years. One notable exception is the Tanquary High area of northern Ellesmere Island, where evidence for repeated fault-controlled uplift until latest Triassic is evident.

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3.3. Phase 3: Passive and cold (Middle–Late Permian) Phase 3 started with a major base-level rise that drowned the previously exposed Sabine Bay and older formations. A deepening-upward succession of calcareous sandstones, and ﬁner clastics, locally replete with brachiopods and bryozoans (Roadian Assistance Formation; up to 200 m thick), recorded a transgression that extended well beyond the pre-existing margins of the basin contemporaneous to one last phase of Permian volcanism (northern Axel Heiberg and west-central Ellesmere Island). Maximum water depth coincided with the widespread deposition of black siliceous shale and spiculitic chert of the van Hauen Formation (up to 600 m), representing basinal to slope deposition (Thorsteinsson, 1974). These mudrocks pass upward into progressively shallower carbonates (lower Degerbo¨ls Formation and glauconitic sandstones (lower Trold Fiord Formation) (Figure 10), Middle Permian (Wordian) carbonates of the Degerbo¨ls Formation contain an even colder biota than their Kungurian and Artinskian counterparts (Reid et al., 2007). Dropstones indicating the presence of, at least seasonal, ice have been reported at that level (Beauchamp, 1994). A major unconformity lies within the Degerbo¨ls and Trold Fiord formations (Beauchamp et al., 2001). The sediments that rest above are deepening-upward cherty carbonates and sandstones that are increasingly fossil-poor with the exception of sponge spicules (Capitanian, upper Degerbo¨ls Formation). Maximum water depth was once again recorded in the widespread development of black shales and cherts. The Lopingian Lindstro¨m Formation (up to 100 m thick) comprises essentially nothing but sponge spicules deposited in a wide range of shallow to deep-water environments (‘‘Glass Ramp’’ of Gates et al., 2004). For reasons not yet fully understood, carbonate factories were eradicated. Cold temperatures certainly played a role, but other factors may have been at play. The Lindstro¨m Formation passes landward into non-fossiliferous, glauconitic sandstones of the upper Trold Fiord Formation. Base level dropped signiﬁcantly near the end of the Permian, subjecting the Lindstro¨m chert to widespread erosion and to contemporaneous progradation of the shallow chert facies well into the basin. The end of the Paleozoic Era was recorded in a deepening-upward wedge of black chert and shale that was deposited in the distal areas and onlapped the shelf margin, while the bulk of the Lindstrom shelf remained subaerially exposed. Filling that depression took the greater part of the Triassic.

3.4. Phase 4: Filling the deep basin (Triassic) As the Paleozoic Era came to a close, the exposed ﬂanks of the Sverdrup Basin were gradually being transgressed. A major base-level rise initiated Phase 4 and the sea reached inland well beyond the previous margins of the basin. A large bathymetric difference, of probably more than 2 km, existed between the basin margins and basin axis. This deep basin began to form in the Early Carboniferous as rifting progressed and continued to expand landwards. It deepened throughout the Late Paleozoic as subsidence rates of the central part of the Sverdrup Basin exceeded long-term sediment supply rates.

Figure 10 basin.

Middle Permian paleogeography with shallow water carbonates, spiculites and sands bordering the deep central

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Figure 11 Late EarlyTriassic paleogeography with major deltas along the eastern and southern margins of the basin that now opened to the southwest.

A new depositional regime was established at the dawn of the Mesozoic Era. The climate was warm and dry, sedimentation was now siliciclastic-dominated, and marine life was sparse and had nothing in common with its Permian counterpart. Major deltas began to prograde into the basin from the south and east (Bjorne Fm) (Figure 11). This huge siliciclastic supply was sourced mainly from Devonian siliciclastic strata that ﬂanked the basin and extended over the craton. The sudden increase in siliciclastic supply suggests regional uplift of the cratonic areas was associated with the tectonic episode that initiated this new phase. Throughout the Early Triassic sandy deltas advanced into the basin (Embry, 1988, 1991a) and pushed the shelf/slope boundary a considerable distance to the north and west. Large, submarine fans were deposited in the deep basin (Blind Fiord Fm) at this time. Substantial sediment was also input from Crockerland (Figure 12). These Early Triassic deltaic and marine strata are up to 2000 m thick (Figure 7) and they represent the ﬁrst major ﬁlling of the central basin. Marginal uplift, followed by a major transgression, occurred at the Early–Middle Triassic juncture and sediment supply was greatly reduced. Dark, bituminous mud and silt (Murray Harbour Fm) was widely deposited and minor progradation occurred. Middle Triassic strata do not exceed 300 m in thickness (Figure 7). Following another episode of uplift and transgression in earliest Late Triassic, siliciclastic supply again greatly increased and a mixed carbonate/siliciclastic shelf (Roche Point Fm) advanced well into the basin. These strata are up to 1,400 m thick. This advance was halted by another transgression in Early Carnian but clastic supply again quickly overwhelmed subsidence and sandy, inner shelf deposits (Pat Bay Fm) built seaward over mud and silt-dominated slope and outer shelf deposits (Hoyle Bay) (650 m). Notably, at this time Crockerland was a major source of sediment and sandy shelf deposits (Pat Bay Fm) extended southward across the entire basin in the west (Embry, 1992). A transgression in Mid–Late Triassic (latest Carnian) again pushed the shorelines to the basin edges. Following this, siliciclastic supply from the east and south greatly increased and during the Norian shallow shelf sands (Romulus Mbr, Heiberg Fm) prograded across prodelta mud and silt (Barrow Fm) (1,000 m) and the basin continued to ﬁll (Embry, 1982). By the end of the Norian the deep water, central portion of the Sverdrup Basin that had existed for over 100 million years was completely ﬁlled and a shallow seaway was present over most of the basin (Figures 7 and 13). During this entire phase of basin development, salt domes and walls, derived from the Otto Fiord Fm, grew upwards in the central portion of the basin and sometimes formed emergent islands (Balkwill, 1978).

3.5. Phase 5: Shallow seas (latest Triassic–earliest Cretaceous) Widespread uplift ushered in the next phase of basin development and in latest Triassic (earliest Rhaetian) much, if not all, of the basin was subaerially exposed. The sea returned in Early Rhaetian but soon major sandy deltas prograded from the east and southeast and much of the eastern and central basin became a vegetated, delta plain

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Figure 12 (A) Facies map for Lower Triassic strata on northern Axel Heiberg Island showing the derivation of siliciclastics from a northern land area (Crockerland). (B) Stratigraphic cross-section of Lower Triassic strata of northern Axel Heiberg Island illustrating three cycles of southward progradation of siliciclastic sediments.

(Fosheim Mbr, Heiberg Fm) under a warm, temperate climate (650 m) (Figure 14). To the west, shoreline and marine shelf sand deposits (MacLean Strait Fm, King Christian Fm) prograded westward over prodelta and shelf mud and silt (Grosvenor Island and Lougheed Island fms) three times between Early Rhaetian and latest Sinemurian (Early Jurassic) (Embry and Johannessen, 1992) (Figure 15). In Early Pliensbachian sediment supply signiﬁcantly waned and the sea transgressed the delta plain leaving a widespread blanket of marine sand in its wake (Remus Mbr, Heiberg Fm) (100 m). From late Early Jurassic (Toarcian) until the beginning of the Late Jurassic sediment supply to the basin was very low and a shallow sea receiving mainly mud and silt (Jameson Bay and McConnell Island fms) occupied the basin (Embry, 1993) (Figure 7). Sandy deposits were mainly restricted to the basin edge and did not prograde very far basinward (Sandy Point and Hiccles Cove fms). The maximum thickness of these strata is 500 m but they are usually less than 300 m over most of the basin. Rift activity related to the initiation of the Amerasia Basin to the northwest continued to expand and began to notably affect the Sverdrup Basin by Middle Jurassic. Narrow rift basins developed along the western margin of the basin in the Prince Patrick Island area. By this time it is assumed that, due to fragmentation by rifting, Crockerland was no longer a source area for the basin. A narrow positive area, the Sverdrup Rim, separated the Sverdrup Basin from the subsiding rift valleys of the proto-Amerasia Basin (Embry, 1992) and it supplied a small amount of sediment to the northwest margin of the basin. Subsidence and sediment supply rates signiﬁcantly increased at the beginning of the Late Jurassic following a brief episode of marginal uplift in the Late Callovian. A series of wave-dominated deltas (Awingak Fm) separated by transgressive events prograded over shelf mud and silt (Ringnes Fm) towards the northwest (Embry, 1993).

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Figure 13 Stratigraphic cross-section of Phase 4 strata of eastern Sverdrup Basin. The deep central basin was ¢lled during this time.

Figure 14 Latest Triassic (Rhaetian) paleogeography with a sandy deltaic plain occupying the eastern and central portions of the basin. Sediment input in the west was low and a shallow sea was present.

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Figure 15 Stratigraphic cross-section of Rhaetian to Sinemurian strata (a 2nd order sequence which forms the lower portion of Phase 5) in the western Sverdrup Basin. The deltaic siliciclastics prograded westward.

Figure 16 Late Jurassic (Kimmeridgian) paleogeography with a broad band of shallow marine sand along the eastern and southern £anks of the basin. A rift shoulder occurred along the northwest margin of the basin that now opened to the southwest.

The deltaic sandstones reach as far as the basin center where this succession is up to 500 m thick (Figures 7 and 16). Also at the beginning of Late Jurassic normal faulting associated with Amerasian rifting extended south of the Sverdrup Basin to the Banks Island area and a seaway that linked the Sverdrup Basin to the Interior Seaway of western North America (see Miall et al., Chapter 9) came into existence. The ﬁnal deposits (250 m) of this phase are shelf muds and silts that occupied much of the basin in latest Jurassic and earliest Cretaceous (Deer Bay Fm) (Figure 7). Shoreline to shallow shelf sands (Awingak Fm) were

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Figure 17 The development of the northwestern £ank of Sverdrup Basin. From Carboniferous to Early Jurassic the basin was £anked to the northwest by a low-lying land area known as Crockerland (A).With the initiation of rifting of the proto-Amerasia basin in Mid-Jurassic, a rift shoulder (Sverdrup Rim) formed the northwest £ank of the basin (B).

deposited on the basin edge during much of this time with highly progradational, delta front sandstones (basal Isachsen Fm) occurring at the top of the succession. The salt structures that had been established in the Late Paleozoic continued to ascend and deform the strata. Large, local variations in thicknesses occur in the vicinity of these structures and a signiﬁcant surge of upward movement occurred at the end of this phase (Jackson and Harrison, 2006). The latest Triassic to earliest Cretaceous interval that was dominated by shallow offshore shelf deposits saw a major change in the general paleogeographic framework of Arctic Canada. This change was due to the ongoing Amerasian rifting which fragmented Crockerland and left a narrow strip of land, Sverdrup Rim, separating the Sverdrup Basin from the proto-Amerasia Basin (Figure 17). The Sverdrup Rim was part of a rift shoulder that marked the entire eastern margin of the Amerasian rift system and served as a minor, intermittent source area for the Sverdrup Basin. When it was emergent, it cut off the long established connection between the Sverdrup Basin and the Chukchi Basin (Hanna Trough) of northern Alaska. A new seaway along the cratonic side of the rift shoulder was formed and it connected the interior seaway of western North America with the Sverdrup Basin. This phase was brought to a close by widespread uplift over the entire basin in the latest Valanginian–earliest Hauterivian interval (Embry, 1991a). This tectonic event in interpreted to coincide with the beginning of seaﬂoor spreading and the creation of oceanic crust in the adjacent Amerasia Basin and the resultant unconformity is sometimes referred to as the ‘‘breakup unconformity’’ (Embry and Dixon, 1994).

3.6. Phase 6: Rejuvenation (Early Cretaceous–earliest Late Cretaceous) Subsidence rates substantially increased in the Sverdrup Basin following the major tectonic episode that initiated this phase. The tectonic regimen of the basin changed from one of slow passive subsidence that had been dominant since Late Permian to one of renewed rifting and extension. Subsidence and sediment supply rates greatly increased and normal faulting and sporadic volcanism also characterize this developmental phase. Thick, coarse-grained ﬂuvial sediments (Isachsen Fm) were the initial deposits in the basin and this depositional regime lasted from Hauterivian to Mid-Aptian (max. 1,000 m). Two transgressive marine intervals punctuate the ﬂuvialdominant succession and the strata progressively onlap the basin margins (Figure 7). Marginal uplift in Mid-Aptian was followed by perhaps the largest transgressive event in the history of the basin. The shorelines were pushed well onto the craton far to the east and south and the Sverdrup Rim was

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Figure 18 Distribution of Cretaceous extrusive and intrusive rocks which represent the on land termination of Alpha Ridge, a hotspot track which crosses the Amerasian Basin.

drowned. The basin received thick deposits of offshore mud and silt (Christopher Fm) which prograded northward (1,100 m). At the end of the Early Cretaceous shoreline and shallow shelf sands (Hassel Fm) prograded into the basin (200 m) (Figure 7) and heralded widespread uplift in latest Early Cretaceous and earliest Late Cretaceous. Thick basaltic volcanic ﬂows (Strand Fiord Fm) of Cenomanian age (Tarduno et al., 1998) were extruded in the basin center during this terminal event and the volcanics thicken northward reaching a maximum thickness of 900 m on northern Axel Heiberg Island (Ricketts et al., 1985). Sporadic episodes of volcanism occurred earlier during this phase in the Axel Heiberg/northern Ellesmere area and diabase sills were intruded into the Sverdrup Basin succession over most of the basin during the Early Cretaceous and earliest Late Cretaceous (Figure 18). This volcanism is a related to the occurrence of a hotspot that lay to the north of the basin at this time (Embry and Osadetz, 1988). The hotspot also was responsible for the construction of a huge volcanic ediﬁce, the 30-km thick Alpha Ridge in the adjacent Amerasia Ocean Basin and subsequent expressions of it may include the thick volcanic ediﬁces of Bafﬁn Bay and Iceland (Forsyth et al., 1986; Lawver and Mueller, 1992).

3.7. Phase 7: Quiescence (Late Cretaceous) Widespread uplift in earliest Late Cretaceous (base Turonian) marked another major tectonic shift and the extension, volcanism and high subsidence and sediment supply rates that characterized that the previous phase ended abruptly. A major transgression initiated this phase and again pushed the shorelines far beyond the margins of the Sverdrup Basin. Following this the basin underwent slow, passive subsidence and for much of Late Cretaceous the basin received mud and silt that contained considerable volcanic ash and bituminous material (Kanguk Fm). The sediment input rate increased in the latter part of the Late Cretaceous and shoreline to shallow marine sandstones prograded into the basin (Expedition Fm) (Figure 7). The maximum thickness of the Late Cretaceous deposits is about 800 m. Widespread uplift marked the termination of this phase and, in the far northeastern portion of the basin, felsic volcanics (Hansen Point Fm) were extruded. These volcanics are thought to be related to the initial rifting of the Eurasian portion of the Arctic Ocean and the associated Morris Jesup– Yermak hotspot (Estrada and Henjes, 2004).

3.8. Phase 8: Fragmentation and uplift (Paleocene–Eocene) Once again the tectonic regime of the Sverdrup Basin underwent a drastic change in latest Cretaceous/Early Paleocene. Following uplift at this time, the basin was transgressed and parts of the eastern portion of the basin underwent rapid subsidence related to compressional loading caused by the opening of the Labrador Sea and Bafﬁn Bay and the consequent impingement of Greenland on the Arctic Islands area in Early Tertiary. Other parts

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of the basin such as the Princess Margaret Arch and Cornwall Arch likely underwent uplift and the adjacent subsiding areas received a high supply of siliciclastic sediment (Ricketts and Stephenson, 1994). The Paleocene– Mid-Eocene succession in these foreland basins consists of a basal transgressive sandstone unit (upper Expedition Fm) followed by regressive marine shale and siltstone (Strand Bay Fm) and a very thick succession of sanddominated, coastal to ﬂuvial deposits with abundant coal (Iceberg Bay Fm) (3,000 m). Deformation progressed through the Paleocene and Eocene and syn-tectonic conglomerates (Buchanan Lake Fm) up to 1,000 m thick were deposited in front of the advancing thrust sheets. The deformation climaxed in Late Eocene (Eurekan Orogeny) and the entire basin was uplifted, bringing to a close a 300 million year history of deposition. The Sverdrup Basin area has remained under compression since that time and has acted as a major source area for the adjacent Amerasia Basin throughout much of the Tertiary. The eastern part of the basin was substantially deformed by thrusting and folding and deformation gradually decreased in a southwestward direction. The western portion of the basin is slightly deformed by broad folds and throughout the basin the salt structures exhibited large vertical growth during this compressive interval.

4. Tectonic Episodes The Sverdrup Basin originated due to an episode of extensional tectonics that affected much of the Caledonian/Ellesmerian Orogenic Belt in Early Carboniferous. The basin was uplifted and deformed in Late Eocene due to compression associated with the opening of the Labrador Sea. During its long history the Sverdrup Basin was affected by numerous tectonic episodes that resulted in widespread unconformities. These in turn allow the succession to be subdivided into large magnitude sequences (Beauchamp et al., 2001; Embry, 1988, 1991a, 1993). The 1st order sequences contain smaller magnitude (2nd and 3rd order) sequence boundaries. The characteristics of these 1st, 2nd and 3rd order unconformities leave no doubt as to their tectonic origin (Embry, 1990, 1997; Beauchamp et al., 2001). The largest magnitude unconformities (ﬁrst order) closely match those recognized in various areas of the North American craton and the origin of these is interpreted to be due to various plate-tectonic processes which affected the continent (Sloss, 1988; Burgess, Chapter 2). It seems reasonable to assume that the smaller magnitude unconformities were also generated by crustal movements driven by plate-tectonic interactions as suggested by Embry (1997). Overall it seems the Sverdrup Basin was affected by a relatively short-lived, tectonic episode once every 5–10 million years and that the magnitude of such an episode varied from quite large (widespread uplift and large changes in tectonic and depositional regimes) to moderate (subdued, marginal uplift and subtle changes in depositional regime).

5. Economic Geology 5.1. Petroleum Exploration for petroleum in the Sverdrup Basin began in 1968 soon after the formation of Panarctic Oils Limited, a consortium of numerous oil companies and the federal government. A string of major gas ﬁelds were discovered between northeastern Melville Island and western Ellef Ringnes Island from 1969 to 1980 (Figure 19). In 1981 a major oil ﬁeld, Cisco, was discovered west of Lougheed Island but following discoveries were disappointing in their small quantities of contained oil. The last well was drilled in 1986 and exploration in the Sverdrup Basin ceased with the collapse of the oil price that year. Strata that are sufﬁciently porous and permeable to act as reservoir rocks are very common in the Sverdrup Basin. In the Late Paleozoic succession potential reservoir units include the Nansen, Canyon Fiord and Sabine Bay fms. The main potential Mesozoic reservoir units are found in the Bjorne, Roche Point, Pat Bay, Heiberg, MacLean Strait, King Christian, Hiccles Cove, Awingak, Isachsen and Hassel formations. The Bjorne sandstones are host to the oil sands of Melville Island and numerous oil and gas shows have been encountered in these sandstones. Most of the gas discovered so far occurs in the MacLean Strait and King Christian formations. These sandstones are of shallow marine origin and have very good to excellent porosity and permeability, especially along the southwestern basin margin (Waylett, 1989). Sandstones of the Remus Member of the Heiberg Fm are also very good reservoirs and contain gas at a few localities. In the basin center area the sandstones of the Heiberg Fm have greatly reduced porosity due to cementation. Large quantities of oil and gas are also in shallow marine sandstones of the Awingak Fm. These reservoir units can have porosities up to 15%. The ﬂuvial sandstones of the Isachsen contain small amounts of hydrocarbons. The sandstones of the Isachsen Fm are very porous and permeable over most of the basin but subsurface occurrences are often in communication with nearby outcrops.

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Hydrocarbon ¢elds, western Sverdrup Basin (from Embry et al., 1991).

Four major petroleum source rock intervals have been identiﬁed in the Sverdrup Basin: the Early Carboniferous Emma Fiord Fm, the Middle Triassic Murray Harbour Fm, the Late Triassic Hoyle Bay Fm and the Late Cretaceous Kanguk Fm (Powell, 1978; Embry et al., 1991). Others may well exist. Notably all the discovered oils appear to have been sourced from the Middle and Late Triassic source strata (Brooks et al., 1992). The Triassic shales contain up to 10% TOC and the organic matter is type II and predominantly of algal origin. These source rocks are mature over much of the western Sverdrup Basin and on the ﬂanks of the basin on the southeast, eastern and northern ﬂanks of the basin (Figure 20). In the basin center the Murray Harbour bituminous shales are overmature due to deep burial and the presence of numerous diabase sills (Embry et al., 1991). The main play in the Sverdrup Basin occurs in Tertiary structures and all the hydrocarbon pools so far discovered occur in such traps. The major facies changes that occur with the Late Paleozoic succession including the occurrence of large reefs suggest that traps with a stratigraphic component way well exist in these strata. The Mesozoic sandstones might also be involved in stratigraphic/structural traps on the ﬂanks of salt structures. Numerous major unconformities occur within the Mesozoic succession and truncation and onlap-related traps as well as incised valley deposits may well be present.

5.2. Coal Bustin and Miall (1991) summarized the coal resources of the basin. Thin seams less than a meter thick occur in Permian, Triassic and Jurassic units. More numerous and thicker beds of coal up to 2 m thick are present in the Early Cretaceous Isachsen Fm. These coals are up to bituminous in rank. The thickest coals are found in the Paleogene–Eocene Iceberg Bay Fm and seams up to 15 m have been recorded. The Paleogene coal seams range in rank from lignite to bituminous.

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Figure 20 Maturity map for the Middle--Late Triassic shales (Murray Harbour and Hoyle Bay formations) which are the main source strata for the discovered hydrocarbons.

6. Summary The Sverdrup Basin is a major rift basin that developed on the Ellesmerian Orogenic Belt in Early Carboniferous. The basin went through eight phases of development. Each phase is characterized by a distinctive tectonic and depositional setting and was initiated by a tectonic episode characterized by widespread marginal uplift and subsequent collapse of the basin. A notable change in climate often accompanied these tectonic changes that may have been a response to somewhat abrupt changes in plate-tectonic spreading direction and/or speed. During the ﬁrst two phases, which occurred during the Carboniferous and Permian, carbonate sediments were deposited in a hot climate. Thick salt deposits ﬁlled the basin early in the development of the basin and huge reefs grew on the shelf margins at various times. The climate became much cooler in Phase 3 (Middle to Late Permian) and spiculitic cherts were deposited on the shelf and in the basin. A new world dawned on the basin with the initiation of the Mesozoic Era and Phase 4. The climate became very warm and siliciclastic sediments poured into the basin from the surrounding cratonic areas. During the Triassic the deep, central basin was gradually ﬁlled and shallow-water sediments reached the basin center by Late Triassic. Phase 5 began with high siliciclastic inﬂux in latest Triassic and Early Jurassic in a warm, humid climate. Siliciclastic inﬂux waned for much of the remaining portion of this phase which extended to earliest Cretaceous. This latter potion of this phase of basin development coincides with the rifting phase of the adjacent Amerasia Basin. Normal faulting related to Amerasian extension affected the southwestern portion of Sverdrup Basin. The basin was rejuvenated in Early Cretaceous (Phase 6) by an extensional episode which coincided with the initiation of seaﬂoor spreading in the adjacent Amerasia Basin. Sediment inﬂux was high and basalts were erupted in the northeast with accompanying widespread diabase dyke and sill intrusion. These igneous rocks represent the terrestrial termination of Alpha Ridge, a hotspot track in the Amerasia Basin. Subsidence and sediment supply signiﬁcantly decreased in earliest Late Cretaceous and bituminous shale was the main deposit of the basin for most of the Late Cretaceous (Phase 7). The basin began to be compressed in the Paleogene due to the opening of the Labrador Sea and the consequent counterclockwise rotation of Greenland. Thick, mainly continental siliciclastics were deposited in local foreland basins (Phase 8) and the entire Sverdrup Basin was uplifted and deformed in Late Paleocene to Eocene. Large gas ﬁelds have been discovered in the western Sverdrup within latest Triassic to Early Jurassic sandstones in Paleogene anticlines. Coal is common in various non-marine units from Permian onward with the thickest seams occurring in the Paleogene strata.

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REFERENCES Balkwill, H. R., 1978, Evolution of Sverdrup Basin. AAPG Bulletin, v. 62, pp. 1004–1028. Beauchamp, B., 1994, Permian climatic cooling in the Canadian Arctic, in Klein, G. D. ed., Pangea: paleoclimate, tectonics and sedimentation during accretion, zenith and break-up of a super-continent, Geological Society of America (Special Paper 288), pp. 229–246. Beauchamp, B., 1995, Permian history of Arctic North America, in Scholle, P. A. and Peryt, T. M. eds., The Permian of northern Pangea, Volume 2: Sedimentary basins and economic resources, Springer-Verlag, Heidelberg, pp. 3–22. Beauchamp, B., and Baud, A., 2002, Growth and demise of Permian biogenic chert along northwest Pangea: evidence for end-Permian collapse of thermohaline circulation. Palaeogeography, Palaeoclimatology, Palaeoecology, v. 184, pp. 37–63. Beauchamp, B., and Desrochers, A., 1997, Permian warm- to very cold carbonates and cherts in northwest Pangea, in James, N. P. and Clarke, J. eds., Cool water carbonates, SEPM (Special Publication 56), Tulsa, USA, pp. 327–347. Beauchamp, B., and Henderson, C. M., 1994, The Lower Permian Raanes, Great Bear Cape and Trappers Cove formations, Sverdrup Basin, Canadian Arctic: stratigraphy and conodont zonation. Bulletin of Canadian Petroleum Geology, v. 42, pp. 562–597. Beauchamp, B., and Olchowy, B., 2003, Early Permian buildups (Tolkien reefs) associated with subaqueous evaporites, Sverdrup Basin, Canadian Arctic: a record of syn- tectonic to post-tectonic reciprocal uplift and subsidence, in Ahr, W. ed., Permo-carboniferous carbonate platforms and reefs, SEPM (Special Publication 78 and AAPG Memoir 83), Tulsa, USA, pp. 133–153. Beauchamp, B., Harrison, J. C., Utting, J., Brent, T. A., and Pinard, S., 2001, Carboniferous and Permian subsurface stratigraphy, Prince Patrick Island, Northwest Territories, Canadian Arctic, Geological Survey of Canada, Bulletin 565, Ottawa, Canada, 93 pp. Brooks, P., Embry, A., Goodarzi, F., and Stewart, R., 1992, Organic geochemistry and biological marker geochemistry of Schei Point group (Triassic) and recovered oils from the Sverdrup Basin (Arctic Islands, Canada). Bulletin of Canadian Petroleum Geology, v. 40, pp. 173–187. Bustin, R. M., and Miall, A. D., 1991, Coal resources, Arctic Islands, in Trettin, H. ed., Innuitian orogen and Arctic platform: Canada and Greenland, Geological Survey of Canada, Geology of Canada, no. 3, Ottawa, Canada, pp. 529–532. Davies, G. R., and Nassichuk, W. W., 1988, An Early Carboniferous (Vise´an) lacustrine oil shale in Canadian Arctic Archipelago. American Association of Petroleum Geologists Bulletin, v. 72, pp. 8–20. Davies, G. R., and Nassichuk, W. W., 1991a, Carboniferous and Permian history of the Sverdrup Basin, Arctic Islands, in Trettin, H. P. ed., Geology of the innuitian orogen and Arctic Platform of Canada and Greenland, Geological Survey of Canada, Geology of Canada, no. 3, Ottawa, Canada, pp. 345–367. Davies, G. R., and Nassichuk, W. W., 1991b, Submarine cements and fabrics in Carboniferous to Lower Permian, reefal, shelf-margin and slope carbonates, Northwestern Ellesmere Island, Canadian Arctic Archipelago, Geological Survey of Canada, Bulletin 399, Ottawa, Canada, 77 pp. Embry, A. F., 1982, The Upper Triassic–Lower Jurassic Heiberg Deltaic complex of the Sverdrup Basin, in Embry, A. and Balkwill, H. eds., Arctic geology and geophysics, CSPG Memoir 8, pp. 189–218. Embry, A. F., 1988, Triassic sea-level changes: evidence from the Canadian Arctic Archipelago, in Wilgus, C., Hastings, B., Kendall, C., Posamentier, H., Ross, C., and Van Wagoner, J. eds., Sea-level changes an integrated approach, SEPM (Special Publication 42), Tulsa, USA, pp. 249–259. Embry, A. F., 1990, A tectonic origin for third-order depositional sequences in extensional basins implications for basin modelling, in Cross, T. ed., Quantitative dynamic stratigraphy, Prentice Hall, Engelwood Cliffs, USA, pp. 491–502. Embry, A. F., 1991a, Mesozoic history of the Arctic Islands, in Trettin, H. ed., Innuitian orogen and Arctic Platform: Canada and Greenland, Geological Survey of Canada, Geology of Canada, no. 3, Ottawa, Canada, pp. 369–433. Embry, A. F., 1991b, Middle-Upper Devonian clastic wedge of the Arctic Islands; Chapter 10, in Trettin, H. P. ed., Geology of the Innuitian orogen and Arctic Platform of Canada and Greenland, Geological Survey of Canada, Geology of Canada, no. 3, Ottawa, Canada, pp. 263–279. Embry, A., 1992, Crockerland – the northern source area for the Sverdrup Basin, Canadian Arctic Archipelago, in Vorren, T., Blackadar, R., Glenister, B., Greiner, H., McLaren, D., McMillan, N., Norris, A., Roots, E., Souther, J., Thorsteinsson, R., and Tozer, T. eds., Arctic geology and petroleum potential, Norwegian Petroleum Society (Special Publication 2), Amsterdam, The Netherlands, pp. 205–216. Embry, A. F., 1993, Transgressive-regressive (T-R) sequence analysis of the Jurassic succession of the Sverdrup Basin, Canadian Arctic Archipelago. Canadian Journal of Earth Sciences, v. 30, pp. 301–320. Embry, A., 1997, The Blind Fiord formation and Blaa Mountain Group (Triassic) of northwestern Axel Heiberg Island, Canadian Arctic Archipelago, in Current research, 1997-B, Geological Survey of Canada, pp. 193–198. Embry, A. F., and Dixon, J., 1994, The age of the Amerasia basin, in Thurston, D. and Fujita, K. eds., 1992 Proceedings International Conference on Arctic margins, MMS94 – 0040, U.S. Department of the Interior, Anchorage, Alaska, pp. 289–294. Embry, A., and Johannessen, E., 1992, T-R sequence stratigraphy, facies analysis and reservoir distribution in the uppermost Triassic and Lower Jurassic succession, western Sverdrup Basin, Arctic Canada, in Vorren, T., Bergsager, E., Dahl-Stamnes, O., Holter, E., Johansen, B., Lie, E., and Lund, T. eds., Arctic geology and petroleum potential, Norwegian Petroleum Society (Special Publication 2), Amsterdam, The Netherlands, pp. 121–146. Embry, A., and Osadetz, K., 1988, Stratigraphy and tectonic signiﬁcance of Cretaceous volcanism, Queen Elizabeth Islands, Arctic Archipelago. Canadian Journal of Earth Sciences, v. 25, pp. 1209–1219. Embry, A., Powell, T., and Mayr, U., 1991, Petroleum resources, Arctic Islands, in Trettin, H., ed., Innuitian Orogen and Arctic Platform: Canada and Greenland, Geological Survey of Canada, Geology of Canada, no. 3, pp. 515–525. Estrada, S., and Henjes, K., 2004, Volcanism in the Canadian High Arctic related to the opening of the Arctic Ocean. Zeitshrift der Deutschen Geologischen Gesellschaft, v. 54, pp. 579–603. Forsyth, D. A., Morel-a-l’Huissier, P., Asudeh, I., and Green, A. G., 1986, Alpha ridge and Iceland: products of the same plume? Journal of Geodynamics, v. 6, pp. 197–214. Fortier, Y., Blackadar, R., Glenister, B., Greiner, H., McLaren, D., McMillan, N., Norris, A., Roots, E., Souther, J., Thorsteinsson, R., and Tozer, T., 1963, Geology of the north-central part of the Arctic Archipelago, Northwest territories (Operation Franklin), Ottawa Canada, Geological Survey of Canada, Memoir 320, 671 pp.

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Gates, L., James, N. P., and Beauchamp, B., 2004, A glass ramp: Shallow-water Permian spiculitic chert sedimentation, Sverdrup Basin, Arctic Canada. Sedimentary Geology, v. 168, pp. 125–147. Harrison, J. C., 1995, Melville Island’s salt based fold belt (Arctic Canada), Geological Survey of Canada, Bulletin 472, Ottawa, Canada, 331 pp. Heckel, P. H., 1986, Sea level curve for Pennsylvanian eustatic marine transgressive–regressive depositional cycles along mid-continent outcrop belt, North America. Geology, v. 14, pp. 330–334. Jackson, M. P. A., and Harrison, J. C., 2006, An allochthonous salt canopy on Axel Heiberg Island, Sverdrup Basin, Arctic Canada. Geology, v. 34, pp. 1045–1048. James, N. P., 1997, The cool-water carbonate depositional realm, in James, N. P. and Clarke, J. A. D. eds., Cool-water carbonates, SEPM (Special Publication 56), pp. 1–22. Lawver, L. A., and Mueller, R., 1992, A hotspot origin for the Canada Basin and the path of the Iceland Hotspot. Eos, Transactions, American Geophysical Union, v. 73(14), 287 pp. Maurel, L. E., 1989, Geometry and evolution of the Tanquary structural high and its effects on the paleogeography of the Sverdrup basin, northern Ellesmere Island, Canadian Arctic, in Current research part G, Geological Survey of Canada, Paper 89-1G, Ottawa, Canada, pp. 177–189. Mayr, U., 1992, Reconnaissance and preliminary interpretation of Upper Devonian to Permian stratigraphy of northeastern Ellesmere Island, Canadian Arctic Archipelago, Geological Survey of Canada, Paper 91-08, Ottawa, Canada, 117 pp. Mickey, M., Byrnes, A., and Haga, H., 2002, Biostratigraphic evidence for the prerift position of the North slope, Alaska and Arctic Islands, Canada, and Sinemurian incipient rifting of Canada basin, in Miller, E., Grantz, A., and Klemperer, S. eds., Tectonic evolution of the Bering Shelf-Chukchi Sea-Arctic margin and adjacent landmasses, GSA (Special Paper 360), Anchorage, USA, pp. 67–75. Nassichuk, W. W., and Christie, R. L., 1969, Upper Paleozoic and Mesozoic stratigraphy in Yelverton pass region, Ellesmere Island, District of Franklin, Geological Survey of Canada, Paper 68-31, Ottawa, Canada, 24 pp. Nassichuk, W. W., and Davies, G. R., 1980, Stratigraphy and sedimentation of the Otto Fiord Formation. Geological Survey of Canada, Bulletin 286, Ottawa, Canada, 87 pp. Patchett, P., Embry, A., Ross, G., Beauchamp, B., Harrison, C., Mayr, U., Isachsen, C., Rosenberg, E., and Spence, G., 2004, Sedimentary Cover of the Canadian Shield through Mesozoic time reﬂected by Nd Isotopic and geochemical results for the Sverdrup Basin, Arctic Canada. Journal of Geology, v. 112, pp. 39–57. Powell, T., 1978, An assessment of the hydrocarbon source potential of the Canadian Arctic Islands, Geological Survey of Canada, Paper 78-12, Ottawa, Canada, 82 pp. Reid, C., James, N. P., Beauchamp, B., and Kyser, T. K., 2007, Faunal turnover and changing oceanography: Late Paleozoic warm-to-cool water carbonates, Sverdrup Basin, Canadian Arctic Archipelago. Sedimentology, v. 249, pp. 128–159. Ricketts, B., 1994, Basin analysis, Eureka Sound Group, Axel Heiberg and Ellesmere Islands, Canadian Arctic Archipelago, Geological Survey of Canada, Memoir 439, Ottawa, Canada, 119 pp. Ricketts, B., and Stephenson, R., 1994, The demise of the Sverdrup Basin: Late Cretaceous–Paleogene sequence stratigraphy and forward modeling. Journal of Sedimentary Research, v. B64, pp. 516–530. Ricketts, B., Osadetz, K., and Embry, A., 1985, Volcanic style in the Strand Fiord Formation (Upper Cretaceous), Axel Heiberg Island, Canadian Arctic Islands. Polar Research, v. 3, pp. 107–122. Sloss, L. L., 1988, Tectonic evolution of the craton in Phanerozoic time, in Sloss, L. L. ed., Sedimentary cover – North American Craton; US., Geological Society of America, The Geology of North America, v. D-2, Boulder, USA, pp. 25–51. Tarduno, J. A., Brinkman, D. B., Renne, P. R., Cottrell, R. D., Scher, H., and Castillo, P., 1998, Evidence for extreme climatic warmth from Late Cretaceous Arctic vertebrates. Science, v. 282, pp. 2241–2244. The´riault, P., 1991, Synrift sedimentation in the Upper Carboniferous Canyon Fiord Formation, SW Ellesmere Island, Canadian Arctic, M.Sc. Thesis, University of Ottawa, Ottawa, Canada, 210 pp. The´riault, P., Beauchamp, B., and Steel, R., 1993, Syntectonic deposition of the Carboniferous Borup Fiord Formation, Northwestern Ellesmere Island, Geological Survey of Canada, Paper 93-1E, pp. 105–112. Thorsteinsson, R., 1974, Carboniferous and Permian stratigraphy of Axel Heiberg Island and western Ellesmere Island, Canadian Arctic Archipelago, Geological Survey of Canada, Bulletin 224, Ottawa, Canada, 115 pp. Trettin, H. P., 1989, The Arctic Islands, in Bally, A. W. and Palmer, A. R. eds., The Geology of North America: the Geology of North America. An overview, The Decade of North American Geology, Boulder, USA, v. A, pp. 349–370. Trettin, H. ed., 1991, Geology of the Innuitian orogen and Arctic Platform of Canada and Greenland, Ottawa, Geological Survey of Canada, Geology of Canada, v. 3, 569 pp. Waylett, D., 1989, Drake Point gas ﬁeld – Canada Arctic Islands, Sverdrup basin, in Beaumont, E. and Foster, N. eds., Treatise of petroleum geology, Atlas of oil and gas ﬁelds, AAPG, Tulsa, USA, pp. 1–27. Ziegler, P., 1988, Evolution of the Arctic-North Atlantic and the Western Tethys. AAPG, Memoir 43, Tulsa, USA, 198 pp.

CHAPTER 14

The Atlantic Margin Basins of North America Andrew D. Miall, Hugh R. Balkwill and Jock McCracken

Contents 1. Introduction 2. The Sedimentary Basins 2.1. Introduction 2.2. Rift basins 2.3. Basins of the southern segment: Bahamas to Newfoundland fracture zone 2.4. Basins of the Grand Banks of Newfoundland 2.5. Basins of the northern segment: Labrador to the Arctic Islands 3. Petroleum Resources 4. Discussion Acknowledgments References

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Abstract The Atlantic margin of North America represents the classic ‘‘Atlantic-type’’ continental margin, notably the margin off the east coast of the United States, which was the site of five deep offshore stratigraphic test holes wells (the Continental Offshore Stratigraphic Test, or COST series) drilled in 1976–1979 on the Georges Bank, the Baltimore Canyon Trough, and the Southeast Georgia Embayment. Data from these holes were used in the development of what have become standard backstripping methods and subsidence models for extensional continental margins. Development of the margin began with the initial rifting of Pangea in the Triassic. Sea-floor spreading began in the central Atlantic Ocean during the early Middle Jurassic, and extended northward past Newfoundland beginning in the Late Jurassic. Active sea-floor spreading generated the Labrador Sea and Baffin Bay between the Cenomanian and the end of the Oligocene. Structural styles vary along the Atlantic margin. Off the southern U.S. margin, the edge of the continental margin was affected by magmatic underplating and extensive volcanism during the Jurassic. The Newfoundland continental margin developed by the processes of crustal thinning and crustal detachment. The Grand Banks area was affected by two distinct phases of rifting and flexural subsidence as extension occurred in the central Atlantic, to the south, from Late Triassic to Early Jurassic, and in the North Atlantic, to the northeast of the bank, from Late Jurassic to Mid-Cretaceous. The thickness of Jurassic-Recent sedimentary deposits on the continental margin locally reaches 25 km. Transects across the margin show a series of largely non-marine rift basins, capped by a breakup unconformity, above which is a seaward-thickening wedge of prograding shallow- to deep-water marine deposits. Evaporites are widespread at the base of this section from the Grand Banks to the Bahamas. Carbonates dominate the remaining deposits in the south, notably in the Bahamas area, but as the North American continent drifted northwestward through the Mesozoic, carbonate sedimentation gradually became less important in more northerly parts of the continental margin. On Georges Bank, carbonate sedimentation ended in the Mid-Cretaceous, whereas on the Grand Banks it had essentially come to an end by the close of the Jurassic. Shallow-marine and deltaic clastics comprise much the remaining succession throughout the length of the Atlantic margin. The discovery of major petroleum resources beneath the Grand Banks in 1979 led to extensive seismic and offshore exploration work there, and additional oil and gas resources have been discovered and developed. Gas reserves have been developed off Nova Scotia, and undeveloped gas reserves are located on the Labrador shelf, but no commercial discoveries have been made in the U.S. portion of the Atlantic margin.

Sedimentary Basins of the World, Volume 5 ISSN 1874-5997, DOI 10.1016/S1874-5997(08)00014-2

r 2008 Elsevier B.V. All rights reserved.

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1. Introduction The modern Atlantic margin of North America extends from northern Baffin Island to the tip of Florida and the southeast corner of the Bahamas, a distance of 7,500 km. This is the largest continuous extensional (‘‘passive’’) margin on Earth (Figure 1), comparable in scale to the extensional margin that encircles the western, southern and eastern margins of the continent of Africa. It includes an important area of petroleum production, the Grand Banks of Newfoundland (Figure 2).

Figure 1 Location of the major basins and tectonic elements of the Atlantic margins of North America. CCF, Cobequid--Chedabucto Fault; CGFZ, Charlie Gibbs Fracture Zone.

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Figure 2

Location map, Grand Banks of Newfoundland (Grant and McAlpine, 1990).

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Figure 3 Plate-tectonic evolution of the Atlantic Ocean, showing the paleogeography at 170 Ma, Bajocian; 160 Ma, Oxfordian; 140 Ma,Valanginian; 120 Ma, Aptian; 100 Ma, Albian; 80 Ma, Campanian; 60 Ma, Middle Paleocene; 40 Ma, Middle Eocene; 20 Ma, Early Miocene. Latitude and longitude grids are shown at 201 intervals. Maps were constructed using the interactive mapping program at http://www.odsn.de/odsn/services/paleomap/paleomap.html

The eastern margin of North America was formed during the breakup of Pangea (Figure 3). The gradual evolution of this breakup, from the Triassic to the present, can be broken down into three broad stages, which suggest a three-fold subdivision of the margin: (1) a southern segment, extending from Florida and the Bahamas to the Newfoundland fracture zone, (2) the Grand Banks area off Newfoundland and (3) the Labrador–Baffin Island margin to the north. Breakup commenced with the development of a rift system in the Triassic that affected the entire North American Atlantic margin and adjacent areas of the flanking continents, from Florida to Nova Scotia, parts of northwest Africa, and most of northwest Europe, from Spain to Denmark (Ziegler, 1988),

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but not Labrador and areas to the north. Sea-floor spreading began in the central Atlantic Ocean, between the Carolina Trough and the Scotian basins (the southern of the three segments referred to above), in the early Middle Jurassic (Figure 3). By 170 Ma (Bajocian), oceanic crust was being generated along the full length of the Central Atlantic spreading system, from the Newfoundland fracture zone to the Bahamas and into the Gulf of Mexico (Klitgord et al., 1988). In the late Middle Jurassic (Callovian), the spreading center shifted outboard, leaving a strip of transitional crust marking what is now the continental margin (Sheridan et al., 1988, and see discussion of the Bahama platform, below). Newfoundland began to separate from Europe (the Iberian Peninsula) during the Late Jurassic (Kimmeridgian, about 153 Ma), but separation was probably not complete until the MidCretaceous (Albian, about 110 Ma). Separation of Eurasia from North America was complete by the end of the Albian, at about 97 Ma (Gradstein et al., 1990; Srivastava and Verhoef, 1992). This Late Jurassic to Albian tectonism affecting the Grand Banks defined the final structure of the central of the three segments of the Atlantic margin. Rifting began in the northern segment, Labrador Sea and Baffin Bay, in the Early Cretaceous (Neocomian). Active sea-floor spreading commenced in the Cenomanian and continued until the end of the Oligocene (95–25 Ma). Magnetic anomaly patterns indicate the presence of a triple-point junction off the southern tip of Greenland between anomalies 24 and 20 (Early to Middle Eocene, 52–43 Ma), after which time Greenland moved with North America (Gradstein et al., 1990; Srivastava and Verhoef, 1992). Spreading ended in Labrador Sea in the Oligocene. A large sediment flux was available to fill the rift basins and extend out onto the continental margin as breakup proceeded. As Ettensohn (Chapter 4, this volume) noted, uplift and erosion of the Appalachian orogen reached a peak as the orogeny came to an end in the Permian, shedding sediment on to the incipient Atlantic margin from Newfoundland to Florida from Triassic time on. It has long been thought that the sediment piles off Labrador and Baffin Island were derived in part by erosion of the Cordilleran mountains of western Canada, with large rivers transporting the detritus eastward across the continental interior during the Cenozoic (McMillan, 1973; Duk-Rodkin and Hughes, 1994). The southern portion of the Atlantic margin, from Florida and the Bahamas to the Newfoundland fracture zone represents a fully developed, and still active ‘‘passive’’ margin with a sedimentary record spanning the MidJurassic to modern, resting on Triassic–Early Jurassic rift basins (Sheridan and Grow, 1988). On the Labrador margin, rifting commenced in the Barremian and extended into the Coniacian (about 130–86 Ma), with post-rift deposition continuing to the present day (Balkwill et al., 1990). Geophysical data indicate a thick sedimentary succession extending along the margin of Baffin Island and into the inter-island channels, including Lancaster Sound (Balkwill et al., 1990). Baffin Bay and the inter-island channels extending westward from it were formed during the approximately 70 Myr period of sea-floor spreading that separated Canada from Greenland. During that interval, Greenland underwent rotation and modest lateral (left-lateral) displacement relative to North America, causing extensional deformation in the southeastern Arctic Island and contractional deformation in the northeast, a phase of tectonism termed the Eurekan orogeny (Trettin, 1991). From Newfoundland to Florida the Atlantic margin includes a broad continental shelf, incorporating four areas of particularly wide shelf underlain by extended continental crust: the Grand Banks off Newfoundland, the Georges Bank, Blake Plateau and the Bahamas Platform (Figure 1). A coastal plain underlain by Jurassic– Cenozoic extensional-margin sediments extends from New York to Florida. Northward from New York the sediment wedge is entirely below sea level, except for a few small areas, including outliers in Atlantic Canada and Labrador, and the fill of Eclipse Trough, on Bylot Island, at the northern end of Baffin Island. Triassic–Early Jurassic rift basins within the Appalachian orogen are exposed along a belt from North Carolina to Nova Scotia. They are commonly called ‘‘Newark basins’’ after the Newark Supergroup, the name applied to the synrift sedimentary succession in the eastern United States (Manspeizer and Cousminer, 1988). The underlying basement consists of the Appalachian orogen from Florida to Newfoundland, and the Precambrian Shield from Labrador to Baffin Island. Remnants of the cratonic Paleozoic carbonate succession are present beneath the younger sedimentary cover off Labrador and southern Baffin Island. The North American Atlantic margin has played an important role in the development of concepts about extensional tectonics and basin development. Seismic-refraction studies were carried out by the Lamont–Doherty Geological Observatory in the 1950s, which led to the first realization of the presence of a thick sediment wedge on the continental margin (Keen and Piper, 1990). Petroleum exploration began in the Scotian basin in 1965. Six JOIDES core holes were drilled in the continental margin off Florida in 1965 (Bunce et al., 1965) and offshore exploration began in the U.S. sector in 1967. Five deep offshore stratigraphic test holes wells (the Continental Offshore Stratigraphic Test, or COST series) were drilled in 1976–1979 on the Georges Bank, the Baltimore Canyon Trough, and the Southeast Georgia Embayment, and provided essential data on the sedimentary environments and burial history of this classic ‘‘Atlantic-type’’ margin. These data were used by Steckler and Watts (1978) and Watts (1981) in their development of the backstripping methodology for the quantitative

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Figure 4 An early numerical model for £exural subsidence on Atlantic-type margins, in which an actual cross-section through the margin of South Carolina is compared to a numerical simulation (Watts, 1981).

investigation of extensional continental margins, and the development of the first numerical flexural subsidence model (Figure 4). Early ideas about application of the ‘‘simple-shear’’ model of crustal extension were applied to the Grand Banks by Tankard and Welsink (1987) and were subsequently developed in detail with the aid of industry and Lithoprobe seismic data. Oil was discovered at Hibernia in 1979, and subsequently this and several other oil and gas fields have been developed on the Grand Banks and near Sable Island, Nova Scotia. Commercial quantities of hydrocarbons have not been discovered in the U.S. sector. The area has now been extensively mapped by the U.S. Geological Survey and the Geological Survey of Canada. There have also been several Deep Sea Drilling Project (DSDP) and offshore drilling project (ODP) wells drilled along this margin. A significant amount of industry seismic and well data has been obtained, particularly within the Canadian portion of the margin, where several commercial petroleum systems have been proven in the Scotian Shelf and Jeanne d’Arc Basin areas.

2. The Sedimentary Basins 2.1. Introduction The basins of the Atlantic margin display the two-phase architecture characteristic of extensional (‘‘passive’’) continental margins, the ‘‘Texas Longhorn’’ or ‘‘Steer’s head’’ configuration of some authors (Dewey, 1982; White and McKenzie, 1988). Rift basins represent the head of the steer, formed by brittle failure during the initial phase of crustal stretching. They are followed by the broad seaward-thickening sediment wedge deposited during the phase of flexural subsidence that accompanies regional cooling and subsidence as the rifted margins

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move away from the sea-floor spreading center (the horns of the steer). These sediment piles characteristically onlap the continental margin over periods of tens of millions of years as the thinned continental crust gradually subsides (Figure 4; Watts, 1981). Rifting was accompanied by the extensive volcanic activity of the Central Atlantic Magmatic Province (McHone, 2000; Hames et al., 2003). Basaltic dikes, sills and flows formed during a 25 Myr period spanning the Triassic–Jurassic boundary over a vast area, including much of west Africa, northern South America, and the Appalachian orogen of North America. The volume of volcanic rocks generated during this period rivals that of other great basaltic accumulations, such as the Deccan Traps of India (Sheridan et al., 1993). These volcanic rocks are an important component of some of the rift-basin fills (e.g., the Palisades Sills of New York), and also occur in the subsurface, such as at the continental margin in the Baltimore Canyon and Carolina Troughs, as noted later. Rift basins indicate the commencement of crustal extension. Formation of the Newark rift basins of the eastern United States began in the Late Triassic and continued into the Early Jurassic. Rifting began later on the North Atlantic margin. The Grand Banks area was affected by two phases of rifting, as a result of the diachronous opening of the Atlantic Ocean, a Triassic phase extending from the south, and a Late Jurassic to Mid-Cretaceous phase extending from the northeast (Tankard and Welsink, 1987). On the Labrador margin, rifting commenced in the Barremian and extended into the Coniacian (Balkwill et al., 1990). In every case, the end of the rifting phase is marked by a breakup unconformity, followed by a seaward-thickening sedimentary embankment, commonly cut by listric growth faults, which developed by detachment within the cover rocks. As is common in many extensional-margin basins that commence development adjacent to a small ocean in tropical settings ( Jurassic Gulf of Mexico, Cretaceous South Atlantic, Cenozoic Red Sea), the first influx of marine waters leads to the deposition of extensive evaporites. The Late Triassic–Early Jurassic Argo and Eurydice evaporite formations, which began to form during the late rifting stage, are widespread in the Grand Banks, Scotian Shelf and Georges Bank areas. Seismic evidence indicates that diapiric evaporites are present beneath all of the U.S. basins, the Baltimore Canyon and Carolina troughs, the Blake plateau basin and the Bahamas basin (Sheridan and Grow, 1988). Evaporites are not present beneath the Labrador Shelf for two reasons: In the MidCretaceous, when oceanic crust began to form off Labrador, this part of the continental margin was at a latitude of about 401N, within a temperate climatic zone not normally associated with evaporite formation. In addition, by the time the Labrador continental margin developed it faced an already large Atlantic Ocean with normal oceanic circulation. North America underwent a rotational northwestward drift (relative to the hotspot frame) during the opening of the Atlantic Ocean (Figure 3; Lawver et al., 2002). In the Late Triassic, when rifting commenced, the Carolina Trough lay at the latitude of the equator. By the beginning of the Cenozoic this basin was at a latitude of about 401 N, and southern Labrador was at a latitude of about 501 N (it is at 521 N at the present day). This steady northward drift can be seen in the change in sedimentary facies, as documented by Jansa (1981). Carbonate sedimentation began in the Mid-Jurassic on the Bahamas platform, and continues there to the present day. On Georges Bank, carbonate sedimentation ended in the Mid-Cretaceous, whereas on the Grand Banks it had essentially come to an end by the close of the Jurassic (Figure 5). The stratigraphy and the structure in each of the North Atlantic Basins exhibit styles of typical passive-margin basins, with many similarities because of their common histories and tectonic genesis (Figures 6 and 7). Synrift parts of the basins are dominated by coast-parallel, linear, largely listric normal faults, linked by sub-orthogonal faults that commonly are expressed seaward as ocean-floor transform faults (Figure 6). In profile, the rift domains are generally asymmetrical, having a major listric normal fault as their border fault. The deep-seated older terranes affect the younger accumulation; for example where the synrift faults of the Scotian Margin and Grand Banks are aligned with and commonly inherited from underlying Paleozoic Appalachian tectonic fabrics. Triassic to Jurassic evaporites, where present, constitute important synrift lithologies and are the generators of structure along the Scotian Shelf and interior parts of the Grand Banks. The continental margin off the Baltimore Canyon and Carolina Troughs is marked by a thick wedge of Jurassic volcanic rocks (Figure 7), which is discussed below in the section dealing with these basins. Throughout the continental margin, synrift structures are overlapped by oceanward-tilted strata of the continental-terrace wedge, consisting dominantly or exclusively of seawardprograding terrigenous clastic sequences. Parts of the Atlantic-margin continental slope along the system have been broken and detached by down-to-basin gravity glide structures. Northern basins have been strongly affected by glacial activity. In the Canadian sector, virtually the entire continental shelf was under ice periodically, with shelf-crossing glaciations commencing in the Late Pliocene at Hudson Strait and in the Mid-Pleistocene on the Scotian margin (Piper, 2005). These events had important consequences for the topmost basin fills, resulting in till tongues, iceberg scour, plume fallout and turbidite deposition, large sediment drifts, catastrophic meltwater releases and widespread sediment failures linked in some cases to glacio-isostasy (Piper, 2005).

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Figure 5 Generalized stratigraphic columns for selected basins along the Atlantic margin, highlighting the diachronous development of carbonate sediments ( Jansa, 1981).

Figure 6 Sketches showing map and cross-sectional views of the classical two-stage model for evolution of passive continental margin of central eastern North America. A, rifting during the Middle Triassic to Early Jurassic; B, Drifting, beginning in the Middle Jurassic, and continuing today (Withjack et al., 1998).

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Northern Jeanned′ Arcbasin

SW

A unconformity

0

A

Grand Banks basin Flemish Pass basin Water bottom

Flemish Cap

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Murre fault zone

10 15 seconds

Central Scotian basin Emerald/Naskapi basins Mohican basin

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Postrift unconformity Fundy fault zone - border fault of Fundy basin and Paleozoic thrust-fault zone

?

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Richmond/Taylorsville basins

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?

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15 seconds

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SE Water bottom

Postrift unconformity Ordovician Taconic suture zone (?) with thrust-fault sequences

Branchville/Jedburg basins

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Jurassic basalts

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transect offset no vertical exaggeration at 4 km/s

Post-rift Salt, syn-rift or post-rift

50 km

Moho

Continental crust Mantle

Moho

Oceanic crust Mantle

Syn-rift

Figure 7 Generalized cross-sections through the Atlantic margin, based on synthesis of deep seismic-re£ection data (Withjack and Schlische, 2005).

2.2. Rift basins The COST-2 well on Georges Banks is the deepest and most important of the offshore wells drilled in the U.S. portion of the Atlantic margin. It was drilled to a depth of 6,667 m, and penetrated a thick section (nearly 2,500 m) of Upper Triassic dolomite, limestone and anhydrite, bottoming in Upper Triassic salt (Manspeizer and Cousminer, 1988). This is the only offshore penetration of a series of some 20 onshore and offshore rift basins that have been mapped south of the Newfoundland fracture zone. Other offshore basins have been delineated on the basis of seismic data. The orientation and basement relationships of these basins indicates that many developed by reactivation of the Late Paleozoic basins formed during the last stages of Appalachian suture. Dextral strike-slip faulting that characterized the last stages of the Acadian orogeny (Gibling, Chapter 6, this volume) was reactivated as leftlateral strike-slip during the Triassic. The Cobequid–Chedabucto Fault is an excellent example. This fault marks the terrane boundary between mainland Laurentia and the Meguma terrane. It became a bounding fault for Late Paleozoic basin development (Gibling, Chapter 6, Figure 16) and evolved into a master fault controlling oblique extension during the Triassic and Jurassic (Withjack et al., 1998). It is now one of the bounding faults of Orpheus Graben. In the U.S. Appalachians, Appalachian thrust faults were reactivated as listric normal faults (Withjack et al., 1998; Figure 8). The transition from rifting to drifting was not synchronous along the length of the southern segment of the continental margin. In the central part of the Florida–Newfoundland segment it commenced in the Middle Jurassic (Withjack et al., 1998). To the south, in the Blake Plateau region, the transition from rifting to drifting appears to have occurred slightly later in the Middle Jurassic (e.g., Klitgord et al., 1988). To the north in the Grand Banks region, the transition occurred much later, in the Early Cretaceous (Srivastava and Tapscott, 1986). Most of the rift basins are asymmetric half grabens, bounded on one side by a system of listric normal faults, and on the other side by sedimentary onlap of the basement (Figure 8). None of the onshore basins exhibit a twosided graben structure. Obliquely crossing normal faults with displacements of up to 3 km, and strike-slip

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Figure 8 Generalized northeast--southwest cross-section through the Newark basin, along the Delaware River. JB, Early Jurassic Boonton Formation; JD, Early Jurassic diabase dikes; JE, Early Jurassic tholeiitic extrusives, with sedimentary interbeds; JPS, Early Jurassic Palisades Sill and related intrusives; BC, border conglomerate; TRP, Passaic Formation (Triassic); TRL, Lockatong formation; TRS, Stockton formation; PC-O, Precambrian and Cambro--Ordvician (Taconic) rocks. Small arrows indicate direction of Taconic and Alleghenian thrust faulting, reactivated as extensional faulting (large arrows) during the Mesozoic (Manspeizer and Cousminer, 1988).

displacements of up to 20 km, cut the basins into rhombic-shaped faults blocks, and attest to the oblique stretching of the Atlantic margin that occurred as Pangea began to break up. The sedimentary succession in these basins may reach 9 km, and consists of a variegated clastic succession of conglomerate, felsic and lithic arenite, siltstone, shale and mudstone, with interbedded basaltic lava flows. Evaporites, eolian sands, coal and kerogen-rich beds are locally important. Siliceous tufas formed locally from hydrothermal systems associated with the lava fields (Birney De Wet and Hubert, 1989). Fossil remains include fish, algae, zooplankton, spores and pollen, the organic remains occurring in sufficient abundance in some cases to qualify the fine-grained deposits as oil shales. Varved and cyclic deposits attest to continually changing climatic conditions, with some workers indicating an orbital control for the cyclicity (Van Houten, 1969; Olsen, 1986, 1990). The North American margin straddled the equator at this time. Tropical temperatures and extreme aridity characterized the environments of northern basins, such as the Fundy Basin, comparable to the modern Death Valley of California. However, progressively wetter facies are present in fills of the Hartford Basin and other basins southwestward along the Atlantic margin, including coals and thicker lake deposits. This regional facies trend may represent a climatic gradient, perhaps coupled with orographic effects and changes in altitude (Olsen, 1990).

2.3. Basins of the southern segment: Bahamas to Newfoundland fracture zone 2.3.1. Introduction to the southern segment The first exploration well was the COST B-2 well drilled in 1975 in the Baltimore Canyon Trough. Up to the end of 1991, 51 additional wells had been drilled. The Baltimore Canyon Trough had 32 exploration wells, the Georges Bank Embayment and the Southeast Georgia Embayment were the location for the remaining 19 wells drilled. Several ODP legs have also explored the regional geology of this area. Hydrocarbon shows have been encountered in several wells, but no commercial quantities have been found. The most significant shows came from the siliciclastic rocks of the Upper Jurassic and Lower Cretaceous interval in the Baltimore Canyon Trough. These prospective zones were drilled by 28 of these wells to test the petroleum system of this area. 2.3.2. South Florida–Bahamas Basin ‘‘The broad, shallow banks and intervening deepwater (800–4,000 m) channels of the Bahama Platform are unlike any other topography on the larger Atlantic margin’’ (Sheridan et al., 1988, p. 329; Figure 9). There has been considerable controversy about the origins of this platform area, since Bullard et al. (1965) pointed out that the best-fit closure of the Atlantic Ocean creates an overlap of the facing continental margins if the Bahama Platform is treated as a continental fragment of North America. A few deep drill holes have been located within the Bahamas. Petroleum exploration reached a depth of 5,700 m in the early 1970s, bottoming in Upper Jurassic carbonates and evaporites, and a few DSDP and ODP holes have penetrated to depths greater than 3 km. The Great Isaac well reached volcaniclastic sediments beneath the Jurassic carbonates. Gravity, magnetic, and deep seismic-reflection data were interpreted by Sheridan et al. (1988) to suggest that the larger, western platform areas (Andros Island, Grand Bahama and other areas approximately to the west of the 781 N meridian) are underlain by transitional continental-oceanic crust formed

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Figure 9

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Remote sensed image of the Bahama Platform. Locations of seismic sections are shown.

during an aborted rifting phase that defined the initial continental margin of North America during the MidJurassic. Later seismic work has showed that this transitional crust is in fact a wedge of Icelandic-type subaerial basalts formed during the first phase of sea-floor spreading (Figure 7; Sheridan et al., 1993). The spreading center moved outboard during the Callovian (about 155 Ma) leaving a strip of thickened early oceanic crust located offshore and parallel to the present continental margin, extending from the Georges Banks area to the tip of Florida, and then offset by right-lateral transform faults into the Gulf of Mexico. The transitional crust, thickened by volcanic activity (discussed below), provided a foundation for carbonate platform development, as this crust subsided and underwent transgression during the Late Jurassic. A broad platform or ‘‘megabank’’ developed from Late Jurassic time on, including the present area of the west Florida shelf, the Florida platform, the Bahama Platform and the Blake Plateau. The present topography of the Bahamas, consisting of flat-topped carbonate platforms cut by deep oceanic channels, is thought to have originated in the Late Cretaceous. The channels have developed over graben, although carbonate aggradation and progradation has obscured, in fact, dramatically altered, the deep structure of the platform (Figures 10 and 11). The presence of shallow-water carbonates as young as Early Cretaceous in the downfaulted blocks beneath the inter-island channels indicates that the channels were not present at the time of deposition. They were formed by faulting during the Late Cretaceous. Sheridan et al. (1988) suggested that this was caused by a phase of oblique, left-lateral northeast-directed contractional tectonism generated by interaction between the Caribbean plate and the North Atlantic plate during the Late Cretaceous. Plate-tectonic reconstructions for the Caribbean region (Pindell and Barrett, 1990) indicate that the continental fragment that constitutes the Greater Antilles (Cuba, Hispaniola) collided with and sutured to the Atlantic plate along the southern margin of the Bahama Platform between the Late Cretaceous and Eocene. Masaferro et al. (2002) described the southwestern Bahama Platform as a foreland basin in front of the northeastverging fold-thrust belt that defines the northern margin of Cuba. Tectonized Bahaman reef carbonate deposits of the Bahama foreland occur adjacent to the Greater Antilles suture zone along the northern coastal belt of Cuba (Ball et al., 1985). The tectonism uplifted and warped the megabank. Uplifted blocks provided the foundations for the present large island and bank areas. Downfaulted or folded areas became the deep, inter-island channels. Seismic data from the southwest corner of Great Bahama Bank demonstrate the development of syntectonic carbonate sedimentation formed during contractional deformation of the edge of the bank during the Miocene (Masaferro et al., 2002).

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Figure 10 A northwest--southeast cross section across the northern part of the Great Bahama Bank. Location is shown in Figure 9 (Eberli and Ginsburg, 1989).

Figure 11 Seismic section across the western edge of the Bimini Bank. Location shown in Figure 9 (Eberli and Ginsburg, 1989).

The present configuration of the continental slopes around the platform and inter-island channels is partly the result of carbonate progradation and partly the result of increased submarine erosion, probably reflecting the gradual acceleration of thermohaline oceanic circulation as the climate became cooler with increasing latitudinal temperature gradients, during the Cenozoic. Eberli and Ginsburg (1989) recognized multiple sequence boundaries in the Cretaceous–Cenozoic carbonate cover (Figure 10), reflecting the interplay between vertical motions of the platform and eustatic sea-level changes. Coral reefs and beaches were formed at 5 m above sea level at 125 ka BP on many of the islands, and constitute the bedrock throughout much of the land areas of the Bahamas and Florida (Neumann and Moore, 1975). Pleistocene and Holocene carbonate eolianites are also important. Two cored holes located on and near the center of the seismic line shown in Figure 11 (within the topsets and foresets of the deposits), drilled into the Bahama platform down to a Middle Miocene level at a maximum depth of 678 m, and revealed much about the stratigraphy, sedimentology, fluids and other features of the platform rocks

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(Eberli et al., 1997). Analyses indicate the importance of ‘‘highstand shedding’’ as the primary control on the depositional evolution of the Neogene western Great Bahama Bank: ‘‘The principal source of sediment to the slope is the extensive offbank transport of suspended, fine-grained bank-top ‘background’ sediment during periods of sea-level highstands when the entire platform was submerged providing the bulk, more than 80%, of the slope sediment. During sea-level falls, the supply of fine-grained sediment to the slope environment is reduced or completely stopped. These deposits of reworked margin-derived material form thin intercalations in the ‘background‘ sediment. Factors controlling the thickness, composition and diagenesis of the deposits, and the formation of discontinuity surfaces are (1) the morphology of the platform, hardgrounds may develop at the base (ramp morphology) or at the top of the lowstand deposits (flat-topped platform), (2) the frequency and amplitude of sea-level changes and (3) the water depth and distance to the margin’’ (Eberli et al., 1997, p. 35). A suite of ODP cores drilled mainly on the lower foresets and bottomsets along the seismic line of Figure 10 provided additional details. They showed that much of the lower foreset deposits consists of carbonate turbidites, while the bottomset deposits include contourites (Betzler et al., 1999). Depositional slopes on the clinoforms reach a maximum of 471, but this is exceptional. Most slopes are less than 41 (Betzler et al., 1999), a point that is obscured by the large vertical exaggerations characteristic of seismic displays. These cores reveal three scales of sequence cyclicity, large scale cycles 60–170 m thick, a medium scale of cyclicity tens of meters thick, and a small-scale cyclicity on a meter scale (Eberli et al., 1997). The highestfrequency sequences represent sea-level cyclicity on a 20-ka time scale, according to Betzler et al. (1999). Three major progradational episodes, of Late Miocene, late Early Pliocene and latest Pliocene age, are considered to indicate sea-level lowstands (Eberli et al., 1997). The major lowstand unconformity dated as Late Pliocene–basal Pleistocene has been interpreted as correlating to the global lowstand that records the build-up of continental ice cover in the northern hemisphere. The modern sedimentary environments and deposits of the Bahama Platform have long been regarded as the type area of carbonate sedimentology. This goes back to the work of such pioneer sedimentologists as Illing (1954) and Beales (1956, 1958). Based on his work in the Devonian Palliser Formation of southwest Alberta, Beales (1958), suggested the term bahamite ‘‘for the granular limestones that closely resemble the present deposits of the interior of the Bahama Banks’’. Systematic study of modern Bahaman sediments by Imbrie and Purdy (1962), Ginsburg et al. (1963), Purdy (1963) and others led to the erection of the classic facies classification of carbonate sediments illustrated in Figure 12.

Figure 12

Modern sediments of the northern Great Bahama Bank. Based on Purdy (1963).

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The modern bank margin is a zone up to 10 km wide covered by skeletals sands and oolite shoals. Detritus composed of fragments of corals, molluscs and foraminifers comprises the commonest particle type. Coral reefs and eolian sands are locally present, particularly on the windward (northeastern) margins of the bank areas. Carbonate eolian dunes may reach heights of 100 m above sea level. Back reef areas, covered by extremely shallow but fully marine waters, are areas of quiet water sedimentation, characterised by peloidal and grapestone muds. Deep-water deposits have been documented by Mullins and Neumann (1979) and Cook and Mullins (1983). Carbonate debris flows and turbidites are common. Deep-water currents along parts of the continental slope, especially on the windward margin, have led to hardground development. Contourite drifts of carbonate sand are common on the adjacent basin floor areas. 2.3.3. Blake Plateau Basin and Carolina Trough The Blake Plateau is named for an area of unusually wide continental shelf located off the east coast of Florida. At the latitude of Cape Canaveral the shelf reaches 460 km in width. The total thickness of the sedimentary cover above the post-rift unconformity reaches more than 13 km (Dillon and Popenoe, 1988). Etheridge et al. (1989) noted an observation made by a number of authors regarding the asymmetry of the continental margin structures bordering the central Atlantic Ocean between the United States and African margins. Large basins on one side face small basins on the other side. Etheridge et al. (1989) suggested that this basic surface evidence can be explained by the development of conjugate extensional margins by a detachment or simple-shear model (Wernicke, 1985). The Blake Plateau was interpreted by Etheridge et al. (1989) as a marginal plateau underlain by continental crust thinned by mid-level crustal detachment. The Carolina and Baltimore Canyon troughs to the north were interpreted as lower-plate margins, with the dip of the master detachment reversing direction (from easterly beneath the Blake Plateau to westerly beneath the Carolina and Baltimore Canyon troughs) at a transfer fault corresponding to the Blake Spur fracture zone (this compares to the interpretation of the Paleozoic continental margin of western Canada based on the detachment model, as summarized in Chapter 5 of this volume; see Figure 7). Deep reflection-seismic data have since become available to test this interpretation (Sheridan et al., 1993), and a different interpretation is now available for the Baltimore Canyon and Carolina Troughs (Figure 7 Sections C and D). Seismic and magnetic data indicate the presence of a seaward-dipping wedge up to 15 km thick at the transition between continental and oceanic crust, interpreted to consist of basalts and volcaniclastic rocks. These rocks are part of the Central Atlantic Magmatic Province, and are thought to have originated as a series of volcanic islands along the length of the U.S. Atlantic margin during the early rift phase of margin development. Beneath this wedge is a zone of interpreted magmatic underplating. Magnetic and gravity data are consistent with this zone consisting of mafic and ultramafic intrusions, probably the source for the overlying volcanic wedge. Seismic data and COST samples indicate that Jurassic rocks at the base of the sedimentary section consist of a mixed terrigenous-carbonate suite. Seismic reflections indicate reef structure at the shelf margin (Figure 13), and salt diapirs in the deep offshore, beneath the foot of the continental slope of the Carolina Trough. A broad carbonate platform covered the area during the Early Cretaceous. Rising sea-levels caused a retrogression of the carbonate platform during the Aptian–Albian, with deeper-water sedimentation replacing a carbonate environment on the outer part of the Plateau (Dillon and Popenoe, 1988). A radical change in depositional style from carbonate-platform to deep-water marls occurred during the Cenomanian over the Blake Plateau Basin. This may have been a result of climatic cooling caused by the northward drift of the continent ( Jansa, 1981) slowing carbonate production to the point that it was unable to keep pace with rising sea levels. By contrast, to the south, over the Bahama Platform, carbonate production continued to the present day. The physiographic boundary between the Bahama Platform and the Blake Plateau is abrupt, marked by an increase in water depth to the north of some 800 m. Dillon and Popenoe (1988) suggested that this does not necessarily indicate a control by tectonism, such as an active fault, but could simply indicate that the climatic gradient across the boundary enabled Bahaman carbonate reefs to ‘‘keep up’’ with increasing accommodation due to sea-level change, whereas those over the Blake Plateau were unable to do so. Landward, the deep-water marls of the Upper Cretaceous succession pass into shallow-water, feldspathic sandstones and claystones. Marine transgression carried shelf sedimentary conditions inland as far as the present fall-line at the foot of the Appalachian Mountains. During the latest Cretaceous (Campanian–Maastrichtian) eustatic sea-levels had risen to flood the entire modern coastal-plain area and all of Florida. A marine channel, the Suwannee Strait, opened a connection between the Gulf of Mexico and the Atlantic Ocean across southern Georgia, dividing the shelf into an area of open-marine carbonate banks and reefs to the south, from an area of sandy, calcareous shales to the north. The channel itself is an area of non-deposition and periodic erosion, as indicated by seismic stratigraphic interpretations (Dillon and Popenoe, 1988).

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Figure 13 Two closely-spaced west--east cross sections through the Blake Plateau Basin, showing the Upper Cretaceous progradation wedge, the thick build-up of Campanian--Maastrichtian strata at the mouth of the Suwanee Channel (left end of Line 19), and the erosional unconformity on the top of the Paleocene. Location shown in Figure 1 (from Pinet and Popenoe, 1985).

The initiation of a long-term drop in global sea levels in the Cenozoic led, in the Late Paleocene or Early Eocene, to a diminishing in importance of the Suwanee Channel and the development of the modern Gulf Stream current, flowing out of the Gulf and northward through the Straits of Florida. Eocene and Early Oligocene deposits, consisting of argillaceous limestones and carbonate muds, built the shelf out to the point where they encountered the erosive effects of the Gulf Stream. Subsequently, during the Late Oligocene and Neogene, generally lower sea levels and cooler climates led to fine-grained clastic aggradation and progradation of the shelf, with many erosional breaks caused by sea-level changes. During periods of low sea level, erosion by the Gulf Stream shaped the continental slope into the Blake Escarpment. The current flowed northward across the shelf, eroding sediment that spilled over the edge of the escarpment, generating turbidity currents that, in turn, eroded gullies and canyons as they moved down the slope. Meanwhile, deep Atlantic circulation at the foot of the slope, flowing southward, built up sediment drifts on the floor of the deep Atlantic Ocean. Dives by submersibles have recorded deep ocean currents of more than 4 km/hr moving along the escarpment at the present.

2.3.4. Baltimore Canyon Trough The Baltimore Canyon Trough lies immediately north of the Carolina Trough, north of Cape Hatteras, between the Norfolk Fracture Zone and the Long Island Platform and Atlantic Fracture Zone (Figure 14). The Baltimore Canyon Trough ranges from 50 to 150 km wide. The trough is separated from the Georges Bank Basin to the north by the broad, low-relief Long Island Platform and separated from the Carolina Trough to the south by the Carolina Platform. The first well, COST B-2 drilled on the United States Continental Margin was drilled here in 1975. There have been 32 additional wells drilled (up to 1991). The geology of this basin is summarized by Grow et al. (1988), and additional analysis of the well and seismic data is provided by Prather (1991). The seaward edge of the trough is bounded by the Mesozoic carbonate and siliciclastic shelf-margin complexes. The Salisbury Embayment to the west of a hinge line defined by the approximate updip edge of the Middle Jurassic strata, contains the younger Upper Cretaceous to Cenozoic (Recent) coastal plain deposits. During the Jurassic the shelf margin prograded oceanward by as much as 40 km. Initial subsidence of the Baltimore Canyon Trough is thought to have been the result of Triassic and Early Jurassic continental rifting and crustal thinning as the North American craton separated from Africa. Rapid but

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Location map, Baltimore Canyon Trough area. Adapted from Grow et al. (1988) and Prather (1991).

variable subsidence along the North American continental margin during the Early and Middle Jurassic was controlled in part by transverse fracture zones, which segmented the margin into sedimentary basins and intervening platforms. The stratigraphy in this basin is similar to the Georges Bank and Scotian Shelf. Dip-oriented seismic lines in the Baltimore Canyon Trough show a wedge of sediment up to 15 km thick overlying a prominent angular unconformity that separates an onlapping sequence of Lower Jurassic and younger sedimentary rocks from underlying Triassic and Lower Jurassic synrift basin-fill and eroded continental basement (Figure 15). Analysis of seismic data and the COST-2 well section indicates that the synrift deposits include sandstone, dolomite, volcanic debris and salt. The basal post-rift deposits comprise limestone, dolomite and anhydrite (Grow et al., 1988). A Jurassic–Cretaceous carbonate platform extends through this area for 650 km, from the Carolina Platform to the Long Island Platform, and constituted the major objective of the drilling program in the 1970s and 1980s (Prather, 1991). The platform complex, varying between 4 and 8 km wide, was situated at the edge of a shelf margin that prograded some 20–60 km during the Middle and Upper Jurassic. This succession was followed, in the Kimmeridgian, by a coastal deltaic and shelf clastic complex. During the Berriasian a raised carbonate rim developed at the edge of the shelf, backed by a clastic belt. An oolite bank developed at the edge of the shelf during the Aptian, and represents the last phase of bank-edge carbonate development. No hydrocarbon accumulations have been located in this basin. Core and sample data indicate a lack of potential source rocks, and there is some indication that hydrocarbons may have dissipated through fractures (Prather, 1991).

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Figure 15 Cross-section through the Baltimore Canyon Trough. Location shown in Figure 14. The section and interpretation are from Grow et al. (1988). Subsequent seismic exploration) has shown that the post-rift unconformity corresponds to the surface labeled ‘‘basement hinge zone’’. The wedge of seaward-dipping re£ectors above this is the initial oceanic crust identi¢ed by Sheridan et al. (1993) (see Figure 7).

2.3.5. Atlantic coastal plain The coastal plain extends from Long Island to Florida (Figure 1), and is underlain by a succession of Middle or Upper Jurassic to Late Cenozoic sediments forming a seaward-thickening wedge. The updip edge of the wedge, the zero isopach, occurs at the Fall Line, along the foot of the Appalachian Mountains. Seaward, the succession thickens to 1,500 m at the New Jersey Coast, to 1,500 m at the Georgia–Florida state line, and to more than 5 km beneath southern Florida. The succession underlying Florida consists mainly of carbonate platform deposits, with some anhydrite and terrigenous deposits at the base of the section (Gohn, 1988). Elsewhere along the coastal plain, the succession is lithologically much more varied, and contains the evidence, in the form of disconformities and unconformities, of the repeated sea-level changes that affected coastal regions during the Cenozoic. Chronostratigraphic study of the successions in outcrop and well records has contributed significantly to the debate regarding the record of global eustatic sea-level cycles (e.g., Miller et al., 2004).

2.3.6. Georges Bank Basin The Georges Bank lies between the southern tip of Nova Scotia and Cape Cod, eastward of the Gulf of Maine Platform (Schlee and Klitgord, 1988; Figure 16). Physiographically, the bank, an important fishing ground, lies at depths of less than 60 m. The main Georges Bank Basin lies beneath the Georges Bank. A deep-water channel separates the bank from the Scotian Shelf. Geologically, this channel lies over the Yarmouth Arch, a prominent Paleozoic basement high, which separates the Georges Bank Basin from the Scotian Basin (Figure 17). Two COST wells were drilled here, and eight exploration wells were drilled in the U.S. portion between 1976 and 1982 (Poppe and Poag, 1993). Oceanic crust offshore is interpreted to be either late Early Jurassic (Toarcian) or possibly Mid-Jurassic (Bajocian/Bathonian), generated during initial separation of Africa and North America. This is the oldest known oceanic crust bordering Canada. Initial spreading is represented stratigraphically by a breakup unconformity in rocks of the continental-terrace wedge. Above that level, Mesozoic–Cenozoic sequences prograded and offlapped progressively seaward as landward- and seaward-tapering wedges, attaining a thickness of 15 km along parts of the outer continental margin. Pre-spreading rifting is interpreted to be Late Triassic–Early Jurassic, based on onshore outcrops of synrift red beds and volcanics in Fundy Basin, and rocks drilled in some of the wells. Basement under the synrift beds consists of the Meguma Group (Lower Paleozoic metasediments of the Meguma Terrane) and Devonian plutons. The bottom 12 m of the COST G-2 well penetrated evaporites thought to represent the Argo Salt (Early Jurassic). The Argo Salt regionally underlies the Iroquois Formation throughout much of the Scotian Basin, but is areally restricted in the Georges Bank Basin. These layers of bedded salt and anhydrite accumulated in isolated elongate depressions and are part of the transition from synrift (continental rifting) to post-rift (drifting) deposition (Poppe and Poag, 1993). Middle and Upper Jurassic strata consist of alternating carbonates (including the Iroquois Formation of Toarcian–Aalenian age) and intervening deltaic clastics (including the Mohican Formation of Toarcian age), with a well-developed shelf-edge carbonate succession straddling the flank of Yarmouth Arch and continuing northeastward from there along a large part of Scotian Shelf (Wade and MacLean, 1990; Poppe and Poag, 1993;

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Figure 16 Location map, Georges Bank and Scotian Shelf basins.

Figure 17 1990).

Cross section from the Georges Bank to the Scotian Shelf. Location is shown in Figure 16 (Wade and MacLean,

Kidston et al., 2002; Figure 16). Interfingering tongues of limestone and shale spanning the Bathonian to Berriasian are referred to several different lithostratigraphic units, the Abenaki, Mic Mac or Mohawk formations. The Iroquois Formation represents a lower part of the broad carbonate platform comprising the Bahama– Grand Banks megaplatform. Siliciclastics dominated in the shoreward part of the basin, where the carbonate platform was bifurcated by a clastic wedge (fed by source terrains in the New England Appalachian highlands) that extended across the central part of the shelf and spread onto the incipient continental slope and rise. The muddy, evaporitic and dolomitic carbonate lithologies and seismic facies of the Iroquois Formation indicate that a barrier reef existed along the seaward edge of the platform during Iroquoian deposition. Behind this reef, the platform is characterized by a series of mounds and oolitic shoals. During Abenaki/Mic Mac time the elevated barrier reef system, which continued to exist along the paleoshelf edge (Figure 16), sheltered the adjacent back reef shelf permitting the accumulation of muddy lithologies. Two broad gaps in the carbonate platform along the southern edge of the basin, however, allowed the siliciclastic deltaic sediments to pass onto the continental slope. Cretaceous strata are seaward-prograding clastic wedges with a regionally developed, seismically discernible thin limestone (‘‘0 Marker’’) at the top of the Barremian. The Cretaceous succession was tilted seaward and the landward part eroded prior to basal Tertiary deposition. Tertiary strata thicken to about 1 km at the present-day shelf edge.

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No wells have been drilled in the Canadian side of Georges Bank; the region is presently under exploration moratorium. Petroleum exploration has not been successful in this basin. No hydrocarbon occurrences have been reported. Source-rock studies of the U.S. wells indicated that the region has poor (richness) potential; the prospectivity of the Canadian segment is perhaps enhanced by the presence of a thicker succession containing salt. 2.3.7. Scotian shelf The continental margin off Nova Scotia (Figure 16) has long been known as the definitive Atlantic-style passive margin: a pull-apart margin followed by thermal sag and a prograding shelf with a carbonate bank, major river delta system and a mobile salt substrate. The three major analogue passive margins are all Atlantic facing, namely the Gulf of Mexico, offshore Brazil and offshore West Central Africa, all which are petroleum productive. The Mesozoic and Cenozoic section of the Scotian Shelf may be divided into four broad tectonostratigraphic units: pre-continental break-up rocks, post continental break-up rocks, deltaic wedge and Mid-Cretaceous and Tertiary transgressive sediments (Wade and MacLean, 1990; Wade et al., 1995; Kidston et al., 2002; Figure 18). In the Late Jurassic, three major deltas existed along the Scotian margin: Laurentian, Sable and Shelburne Deltas. Carbonate banks, ramps and reefal complexes flourished on stable platforms and interdeltaic regions. By Early Cretaceous, carbonate deposition ceased and the Sable Delta became the dominant depositional system in the region and expanding during a period of relative sea level highstand. It may be more than 5 km thick in the Sable sub-basin. Small outcrops of Early Cretaceous sediments onshore in Nova Scotia and New Brunswick represent the drainage systems — some developed over karst terrain — that fed the delta (Stea and Pullan, 2001; Gobeil et al., 2006; Falcon-Lang et al., 2007). In Nova Scotia, thick (up to 30 m) saprolites developed on deeply weathered granitic rocks as a result of warm latest Paleozoic and Mesozoic climates, and such widespread regoliths may have fed silica sand and kaolinite-rich clays to Cretaceous drainages (O’Beirne-Ryan and Zentilli, 2003; Pe-Piper et al., 2005). By Middle Oligocene, a major lowstand exposed the entire shelf. A series of shelf-margin deltas and upper slope canyon systems developed, providing sources and conduits for coarse-grained clastic sediments to reach deepwater depocenters. The deepwater Scotian Slope is located on the seaward portion of the 25+ km thick Mesozoic and Cenozoic sedimentary prism that was deposited along the rifted continental-oceanic crustal hinge line zone. Early synrift Late Triassic–Early Jurassic sediments and evaporites (salts) were deposited in a heavily faulted and rifted terrane. During the subsequent drift phase that followed the separation of Morocco and Nova Scotia, the shelf prograded seaward, with the slope region the locus for deposition of fine-grained sediments. Shelf advancement was punctuated by periodic sea level falls with resultant gravity slides and turbidite flows carrying coarser-grained sediments into very deepwater, with deposition over and around the seafloor topography created by salt halokinesis. Salt structures of the Slope Diapiric Province are widespread, and include allochthonous salt canopies that were mobilized as a result of loading by prograding slope sediments, migrating up to 100 km basinward and causing the overlying strata to become detached and extended (Shimeld, 2004; Ings and Shimeld, 2006). The slope area has been significantly modified by subaerial and submarine erosion during lowstands, especially in the Tertiary and even quite recently, with major canyons carved into the slope following Pleistocene glaciation. Six erosional unconformities ranging in age from Oligocene to Pliocene have been identified on a seismic section across the Laurentian Channel, indicating that this channel has a long history as a major drainage conduit from the continental interior (Wade and MacLean, 1990). As noted above, the Scotian margin represents a classic ‘‘Atlantic-type’’ continental margin. Cross sections interpreted from seismic-reflection data (Figure 19) reveal the presence of a suite of rift basins at depth. The breakup unconformity appears nearly everywhere to be blanketed by the Argo Salt, which has also been deformed into a series of diapirs. However, the Argo has been dated as pre-breakup in age (Hettangian). The seismic interpretation shown in Figure 19 suggests a smearing out of the salt along the unconformity, a result of salt mobilization during loading. However, structural details in evaporites are difficult to resolve on seismic data and the interpretation shown in this cross-section is probably over simplistic. The Mesozoic–Cenozoic stratigraphic succession which follows the evaporites consists of a series of seaward-thickening clinoform packages, cut by numerous listric faults. The thermal evolution of the upland region bordering the Scotian Basin probably represents in part cooling due to rift-flank uplift and exhumation, along with changes in surface temperature regimes (Grist and Zentilli, 2003). The basin was affected by tectonic events through the Cenozoic, as a result of motion on the linked strikeslip fault systems of the Newfoundland Fracture Zone, southwest Grand Banks Transform and Cobequid– Chedabucto Fault Zone (Pe-Piper and Piper, 2004). Local magmatic rocks formed during Triassic to Early Jurassic rift phases and Early to Mid-Jurassic post-rift phases (Pe-Piper et al., 1992). Close to the Newfoundland Fracture Zone lie the Fogo Seamounts of Cretaceous age, which represent local volcanic activity associated with detachment faulting and the separation of Iberia and the Grand Banks (Pe-Piper et al., 2007). Large failures of the

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Stratigraphic table for the Scotian Basin (Kidston et al., 2002).

continental margin have become a prominent feature of the margin in the past 40 million years, with the largest failures taking place during non-glacial periods but frequent smaller failures during periods of deglaciation (Campbell et al., 2004). Hydrocarbon source rocks range from Middle Jurassic (Misaine Shale) to Albian (Sable Shale). The Upper Jurassic to Lower Cretaceous Verrill Canyon Shales are the most likely sources for gas, condensate and oil that have been discovered in the Mic Mac, Missisauga, Logan Canyon and Dawson Canyon sandstones. Most of these hydrocarbon discoveries are within shallow-water deltaic-wedge clastics.

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Cross-section across the Scotian Shelf and slope. Location is shown in Figure 16 (Wade and MacLean, 1990).

2.4. Basins of the Grand Banks of Newfoundland The Grand Banks of Newfoundland became an important Canadian petroleum province with the discovery of oil at the Hibernia P-15 well in 1979 (Figure 20). Several other fields have been developed since that time. As a result of this exploration activity, a wealth of drilling and seismic data has been collected from the area, and several specialized collections of research papers have been published containing descriptions of the geology of the Grand Banks, and of other comparable Atlantic-margin areas (e.g., Beaumont and Tankard, 1987; Tankard and Balkwill, 1989). There is, in addition, a major Decade of North American Geology chapter discussing this area (Grant and McAlpine, 1990). The key to understanding the regional development of the Grand Banks is to explore the relationship between the evolution of the oceanic crust off the Grand Banks, and the orientation, sense of movement and timing of the complex pattern of extensional structures that dominate the geology of the area (Tankard and Welsink, 1987; Welsink et al., 1989; Figure 20). Most published discussions deal with the Jeanne d’Arc basin, which is the main focus of this section. There were four broad phases of basin subsidence. (1) The first phase, of Late Triassic–Early Jurassic age, lasted some 20–30 Myr, and resulted in the deposition of a red bed complex capped by evaporites. (2) Movement on major northwest–southeast faults, including the Newfoundland fracture zone and the parallel transfer faults on the Grand Banks initiated the fragmentation of the shelf platform. (3) A phase of slow thermal subsidence during the Early and Middle Jurassic led to the deposition of a monotonous mudstone–carbonate succession (epeiric basin phase in Figure 21). (4) Africa began to separate from Nova Scotia in the Middle Jurassic, at about 175 Ma, and this led to about 40 km of continent–continent displacement along the Newfoundland fracture zone and the transfer faults until Valanginian time. Synrift deposits developed in the Jeanne d’Arc and other basins. This phase of movement came to end with the separation of Iberia at about 115 Ma. (5) During the post-rift phase, which followed, oceanic crust began to develop off the northeast margin of the Grand Banks, and crustal extension led to the development of a suite of crossing faults oriented northeast–southwest (Figure 20). The stratigraphic section of the Jeanne d’Arc basin is illustrated in Figure 21 and the structural geology is shown in two seismic cross sections, Figures 22 and 23. The evolution of the structure was complex. During the Late Jurassic and Early Cretaceous, extension oriented northwest–southeast initiated the major basin-bounding master faults, including the Murre fault, and caused right-lateral strike-slip displacement on a suite of faults perpendicular to the basin-bounding master faults (Tankard et al., 1989). The Murre Faults is interpreted as a master detachment, flattening out at a depth of about 22 km (Figure 22). During the Late Cretaceous, the changing trajectories of sea-floor spreading that led to the opening of the North Atlantic Ocean, imparted extensional stresses on the Grand Banks in a direction more or less perpendicular to the earlier pattern, that is, northeast–southwest. The master faults display evidence of right-lateral strike slip superimposed on the earlier extensional movement, while the northwest–southeast-oriented strike slip faults developed extensional dip slip

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Figure 20 Structural framework of the Grand Banks, showing major faults, magnetic lineaments and the trajectories of the developing oceanic crust (dashed lines). The parallelism between structural trends in the continental and oceanic crust indicates a close genetic linkage (Welsink et al., 1989).

during this phase. This was the major phase of movement that developed the Orphan Basin, on the northeast flank of the Grand Banks, facing the Late Cretaceous oceanic crust of the North Atlantic Ocean. Extension of the crust that created Jeanne d’Arc Basin has been calculated to be about 50%, compared to a regional average of about 20%. The basin is 17 km deep, including a 3-km-thick post-rift succession (Tankard et al., 1989). Each of the phases of subsidence and deformation can be linked to discrete phases of sedimentation in the basin. The earliest phase of subsidence, during the Late Triassic to Early Jurassic (Carnian–Sinemurian) was accompanied by deposition of a suite of non-marine argillaceous red beds and evaporites, the latter deposited primarily in very shallow-water to non-marine sabkha environments. A carbonate-dominated lagoon developed during transgression in the Pliensbachian, and this was followed by the uniform subsidence of the post-rift epeiric basin phase (Toarcian–Callovian; Figure 21) during which calcareous shales and limestones were deposited. Phase two subsidence (Callovian–Aptian) created accommodation for six unconformity-bounded sequences, each spanning 7–10 Myr. It is this synrift succession that is involved in the large roll-over structure in the hanging wall of the Murre Fault (Figure 23), and the structure has been further modified by salt diapirism. Folds, minor internal faults and erosional unconformities yield a detailed history of subsidence during this phase, the details of which are beyond the scope of this chapter. The Hibernia oilfield contains multiple intervals of oil-bearing reservoir rock, including the Ben Nevis and Avalon Formations, the Catalina Member, and the Hibernia and Jeanne d’Arc Formations. Most of the resources, however, are contained in the Ben Nevis–Avalon and Hibernia reservoirs. The petroleum was sourced from the Egret Shale, of Kimmeridgian age, the same age as one of the principle source rocks in the North Sea Basin. A widespread unconformity at the Barremian–Aptian contact marks the end of rift phase two. Subsequently, deposition took place more uniformly across the Grand Banks during the flexural subsidence phase.

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Figure 21 Stratigraphic table for the Jeanne d’Arc basin, showing the unconformity-bounded sequences, tectonic phases and the evolving style of basin development (Tankard et al., 1989).

A continental-terrace wedge was developed, showing a gentle northward tilt. Late Cretaceous to Eocene successions include shelf-slope sand bodies on the western basin margin and small, sand-prone submarine fans on the basin floor that were fed by canyons incised through the western margin (Deptuck et al., 2003). The Orphan Basin on the northeast margin of the Grand Banks exhibits a similar structural and stratigraphic style to the Jeanne d’Arc Basin, with the phase 2 subsidence the dominant basin forming episode. Well control in the Basin is sparse, with seven exploration wells drilled in the basin up to 2003. The fist well was drilled on a basement high, but encountered only thin Mesozoic sediments before entering the Paleozoic. A DSDP well drilled on the Orphan Knoll has also provided some geologic data (Smee et al., 2003). Basins of the Grand Banks and Iberian facing margins originated as intra-continental rifts during the early stages of separation of North America from Europe. In consequence, they share many stratigraphic features in

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Figure 22 A northwest--southeast oriented Lithoprobe seismic line through the southern end of the Jeanne d’Arc Basin, showing the half-graben form of the basin, the marginal roll-over structure, and the basal detachment, which is interpreted to sole out at a depth of about 22 km along a band of subhorizontal re£ectors. Location is shown in Figure 20 (Tankard et al., 1989).

common, although local paleogeographic settings resulted in systematic variations in storm and tidal energy that influenced the nature of their reservoir sand bodies (Hiscott et al., 1990a, 1990b).

2.5. Basins of the northern segment: Labrador to the Arctic Islands Labrador Sea, Davis Strait, Baffin Bay and Nares Strait collectively comprise a 3,500 km-long seaway, initiated by Cretaceous rifting and plate separation, connecting the Atlantic and Arctic oceans. In simplistic terms, this vast tectonic domain consists of elongate, northwest-trending, stretched and deeply subsided, small, rhombic ocean basins, which are linked by narrow, northeast-trending transform zones. Plate motion along this lithospheric network was initiated in Early Cretaceous (?Neocomian) (or possibly Late Jurassic) and ceased in latest Eocene or Early Oligocene (prior to anomaly 13) when ocean spreading aborted and shifted to its present-day location between Greenland and Europe (Balkwill et al., 1990). Twenty-eight exploration wells were drilled on the Labrador Shelf and three holes were drilled off southern Baffin Island during the period 1974–1983. There has been no drilling activity on the Canadian side of Baffin Bay or further north, whereas there has been exploration on the Greenland side of the bay. The Canadian plate margin of this tectonic region has first-order attributes typical of Atlantic-type passive margins: a nearly flat continental shelf margin merges outboard with a more steeply inclined continental slope; the inner part of the shelf is an erosional surface developed on Precambrian crystalline basement (and locally, on Lower Paleozoic platform strata); the outer shelf and slope are constructed by a thick prism of seaward-dipping Cretaceous and Tertiary terrigenous clastics; the more landward part of the shelf prism lies on extended cratonic crust, locally containing large fault-bounded wedges of Lower Cretaceous and lower Upper Cretaceous synrift clastics; the outer part of the prism lies on and is intercalated with thick Upper Cretaceous and Tertiary basalts, representing transitional cratonic/oceanic crust or oceanic crust (Balkwill et al., 1990). 2.5.1. Labrador shelf Initial Labrador Basin rifting is probably recorded by Early Cretaceous Alexis Volcanics (K-Ar age 122 Ma at Bjarni H-81) (Figures 24 and 25). Magnetic anomaly 33, of Santonian age, is the oldest recognized in the Labrador Sea. Anomaly 27 (61 Ma, Early Paleocene) indicates the onset of spreading in the northern Labrador Sea (Harrison et al., 1999). The presence of Late Jurassic mafic dikes along the West Greenland coast indicates the possibility of older synrift rocks in undrilled structures. Bjarni Formation synrift strata are divisible into lower and upper members. The former is dominated by arkosic, coaly, generally coarse-grained fluvial sandstones and the latter by siltstones and sandy shales. The transition from rifting to sea-floor spreading and flexural subsidence of

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Figure 23 Seismic line oriented northwest--southeast through the Hibernia ¢eld. The ¢eld is a roll-over trap adjacent to the Murre Fault, at left. Salt that has risen along the Murre fault is shown by the pattern of squares. Location is shown in Figure 20 (Tankard et al., 1989).

the continental margin lies in the lower part of the thick, marine, shale-dominated Markland Formation. The remainder of the shelf prism deposited during this phase is shale dominated, with some locally developed, inboard sandstone intervals in Upper Cretaceous and Lower Tertiary (Freydis, Gudrid and Leif members). The termination of sea-floor spreading is represented by the regionally mappable Kenamu–Mokami sequence boundary. The upper part of the terrace prism (Mokami and Saglek formations) is increasingly coarser upward, having been deposited as the Labrador (and Baffin) coasts were uplifted and eroded as a result of post-rift lateral heat flow toward the craton, amplified by flexural uplift from sediment loading of the seafloor. Hopedale and Saglek sub-basin depocenters are separated by a broad basement arch (Okak Arch). Synrift faults in Hopedale Basin occupy a relatively wide part of the shelf; they consist of semi-orthogonal patterns of seaward-dipping basement faults, linked by short transfer faults (Figures 24 and 25). Lying discordantly above the

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Figure 24

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Hopedale Basin, Labrador Shelf, showing depth to basement and well locations (Balkwill et al., 1990).

basement fault array in the outer part of the shelf, are clusters of seaward-dipping growth faults, detached in Markland shales, and having rotated structural traps at mid-Tertiary levels. In contrast with the foregoing, large synrift faults in Saglek Basin are landward-dipping, leading to the likelihood that Okak Arch forms a lithospheric zone across which there was a polarity reversal in crustal delamination. Onshore in Labrador, a widespread Mesozoic regolith was probably present, the erosion of which generated economic, iron-rich lacustrine deposits of the Late Cretaceous Redmond Formation (Umpleby, 1979). Gas and minor oil shows were discovered at the Bjarni well off southern Labrador, and condensate in the Hekla well in the southeast Baffin shelf. Five large gas fields were subsequently discovered in synrift structures in Hopedale Basin (Bell and Campbell, 1990). Synrift, Lower Cretaceous Bjarni clastics, forming the reservoir level at Bjarni, North Bjarni, Hopedale and Snorri, were probably sourced by interbedded coals and carbonaceous shales. Lower Paleozoic dolomite is the reservoir at Gudrid.

2.5.2. Davis Strait transform Reaching northeastward from the northern end of Saglek Basin, Davis Strait Transform is a broad zone of lithospheric, left-hand strain translation linking Labrador Sea and Baffin Bay (Balkwill et al., 1990). The transform zone consists of transtensional and transpressional basement structures (positive and negative flowers),

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Figure 25 Seismic cross-section through the Hopedale Basin, Labrador Shelf. Approximate location shown in Figure 24 (Balkwill et al., 1990).

with draped, rotated and faulted Cretaceous and Tertiary sedimentary and volcanic cover rocks. Truncation/ onlap relationships indicate that volcanism and mafic intrusion ended with cessation of Labrador Sea spreading. Gas and condensate were recovered from Paleocene (Gudrid) sandstones, draped over a volcanic flow at the Hekja structure, in the southwestern part of the transform zone (Bell and Campbell, 1990). Six dry holes have been drilled on the West Greenland shelf, in structures approximately at the northeastern end of Davis Strait Transform.

2.5.3. Baffin Bay No exploration wells have been drilled along the Baffin Island shelf or slope. Seismic control is limited to the northern part of the shelf, near the mouth of Lancaster Sound. The ages and nature of the Mesozoic–Cenozoic stratigraphic succession have been interpreted from outcrops on Bylot Island and vicinity (Eclipse Trough: Miall et al., 1980; Harrison et al., 1999), and it is considered probable that the tectonic history of the shelf can be anticipated from events and stratigraphy known along Labrador Shelf, Davis Strait and the West Greenland shelf. The succession, which is as thick as 14 km at the mouth of Lancaster Sound, is interpreted to consist of Lower Cretaceous (Bjarni-like) synrift, largely non-marine clastics, overlain by Upper Cretaceous–Lower Paleocene marine shales, and Upper Paleocene and younger marine shales and sandstones. The interpretation carries a caveat of risk, because of uncertainty with regard to the age and kinematics of Baffin Bay tectonism. The northern Baffin shelf has (at times) attracted exploration interest because of the presence of large structures, the probability of sandstone reservoirs, the presence of an active oil seep at Scott Inlet (at a midpoint along the Baffin Island coast), and outcrops of organic-rich Upper Cretaceous marine shales (Kanguk Formation) on Bylot Island. Lancaster Sound, a narrow, fault-bounded submarine graben, is an appendage to the Mesozoic–Cenozoic succession in northern Baffin Bay, and is presumed to hold similar stratigraphy. It contains 7 km of Paleozoic and younger sedimentary rocks. As reported by Balkwill et al. (1990, p. 333) it has been suggested that this graben represents a Precambrian aulacogen reactivated in the Mesozoic. The present structure of the basin is probably the result of extensional tectonism associated with the rotation of Greenland away from Baffin Island.

3. Petroleum Resources There are four areas along the Atlantic Margin where commercial and ‘‘stranded’’ (not-commercial-at-thistime) hydrocarbons have been found: Sable Island area on the Scotian Shelf, Jeanne d’Arc Basin on the Grand Banks, the Hopedale Basin on the Labrador Shelf and the Saglek Basin on the Baffin Shelf.

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The first commercial production from the Scotian Shelf Basin was 7.1 106 m3 (44.5 million barrels) of oil from Panuke and Cohasset in the Sable Island area between 1992 and 1999. Gas from the Sable Basin has since been under production since 1999 and has produced 31 109 m3 (1.1 trillion cubic feet) from Thebaud, Venture, North Triumph, Alma and South Venture. It is expected that the total recovery from this project alone could be 85 109 m3 (3 tcf ) of gas. (Canada–Nova Scotia Offshore Petroleum Board, Annual Report 2006–2007). Another project, Deep Panuke is being planned. The Canada–Nova Scotia Offshore Petroleum Board has published that the ultimate recoverable resource in this area could be 159 109 m3 (5.6 tcf ) of gas and 28.8 106 m3 (181 million barrels) of oil/condensate from 26 discoveries (Canada–Nova Scotia Offshore Petroleum Board, 2000). Wade et al. (1989) reported that this Scotian Shelf area including the Shelbourne, Abenaki, Mohican and Sable Sub-basins could have the potential of 510 109 m3 (18 tcf ) of gas and 170 106 m3 (1.07 billion barrels) of oil/condensate. A similar study interpreted that the Laurentian Sub-basin, the easternmost extension of the Scotian Basin, could have the potential of 227–255 109 m3 (8–9 tcf ) of gas and 95–110 106 m3 (600–700 million barrels) of oil (MacLean and Wade, 1992). There also has been further additional assessments of the deep-water gas potential on the extensive Scotian slope (Kidston et al., 2002). The Jeanne d’Arc Basin has been under production since 1997 and has produced 116 106 m3 (733 million barrels) of oil from Hibernia, Terra Nova and White Rose. This prolific basin has the potential of an additional 377 106 m3 (2.37 billion barrels) of oil and natural gas liquids and almost 170 109 m3 (6 tcf ) of gas from the above 3 fields and 15 other discovered accumulations (The Canada–Newfoundland and Labrador Annual Report, 2006–2007) There has been some recently announced expansions on the existing fields as well as the probable production from the Hebron discovery. Bell and Campbell (1990) estimated that the potential of the Jeanne d’Arc Basin and area has the approximate potential of 1335 106 m3 (8.4 billion barrels) of oil and 340 109 m3 (12 tcf ) of gas. Oil exploration has been expanding outside of this prolific area into the Flemish Pass and Orphan Basin in the last few years. On the Labrador Shelf the resources estimated for the five gas discoveries of North Bjarni, Gudrid, Bjarni, Hopedale and Snorri total 119 109 m3 (4.2 tcf ) and 20 106 m3 (123 million barrels) of natural gas liquids. The GSC estimate from Bell and Campbell (1990) for the resource potential of the Labrador to Baffin Shelf could be from 80–134 106 m3 (500–843 million barrels) of condensate and 539–737 109 m3 (19–26 tcf ) of gas.

4. Discussion The outlines of the continents were what first caught Wegener’s attention in the 1920s, and inspired him to draw his famous map showing how the world’s continents might once have fit together (Wegener, 1929). From the outset, his map contained problems, notably the fit of Greenland back against Canada, which opened up a hole between Greenland and the Canadian Arctic Islands. This was the beginning of the notorious (to Danish and Canadian geologists) ‘‘Nares Strait problem’’ (Dawes and Kerr, 1982). There were other overlaps, too, such as that in the vicinity of the Bahama Platform, when Africa was pushed back against the United States Atlantic margin (referred to earlier in this chapter; see Bullard et al., 1965). Modern research in extensional tectonics has provided answers to these problems, and the Atlantic margin of North America is where many key solutions have been developed. From the very beginnings of the modern era of plate tectonics, the Atlantic margin has provided a model for comparison. The western Paleozoic margin of North America was compared with the modern Atlantic margin in the early 1970s (see Chapter 5). Detailed studies of paleomagnetic lineations in the Atlantic Ocean have yielded extremely well-constrained kinematic models of Atlantic spreading (e.g., Srivastava and Tapscott, 1986; Srivastava and Verhoef, 1992), providing, for example, relative rotation paths for Greenland and North America which need to be accommodated to the observable geology along the contact between these plates in the northeastern Arctic Islands. This has now largely been done (see discussion in Harrison et al., 1999; the Nares Strait problem has largely been solved). Detailed studies of crustal extension, aided by the COCORP and Lithoprobe projects (see Chapter 17) have provided explanations for what happens when continental crust is stretched and breaks, with the generation of new oceanic crust along the fracture. Stretching of the lower crust, the pure shear model of McKenzie (1978), could help to explain how continental margins could be extended to positions that overlap if a simple geometric fit of post-rift margins was attempted. However, the simple-shear model developed initially by Wernicke (1985) has become the basis for a much more important and comprehensive explanation of Atlantic margin evolution. The collection of research papers on the North Atlantic margins edited by Tankard and Balkwill (1989) is a landmark in the field of applied plate tectonics, because it represents a rigorous, detailed examination of the simple-shear model as applied to a single, complex, extensional margin. The constituent papers demonstrate, through detailed analysis of many of the basins bordering the North Atlantic Ocean, how

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brittle failure of the upper crust is translated into extension along detachments that may bottom out at depths of more than 20 km, thereby explaining substantial modifications to the configurations of the continental margins. This analysis provides explanations for the complex structures in the resulting basins that are of so much interest to the petroleum geologist. The development of the Atlantic margin of North America reflects a relatively uncomplicated evolution of the Atlantic spreading center, which, while repeatedly undergoing subtle changes in trajectory, did not involve major jumps in the position of the spreading center, unlike the evolution of the European margin (Ziegler, 1988), and unlike the evolution of Tethys by the repeated fragmentation of the northern Gondwana margin (AudleyCharles and Hallam, 1988). For all these reasons, the emergence of the term ‘‘Atlantic-type margin’’ is entirely understandable.

ACKNOWLEDGMENTS Thanks are due to Martha Withjack and R. W Schlische for arranging for us to receive a copy of their recent seismic interpretations of the U.S. portion of the Atlantic margin (Figure 7), and to Martha and to Bob Sheridan for their comments on the structural interpretations. David Piper and Marcos Zentilli provided many useful references relating to recent work, at the request of Martin Gibling, who reviewed the manuscript and provided many important corrections and additional information.

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A. eds., The Atlantic continental margin: U.S., The Geology of North America, Geological Society of America, Boulder, CO, v. I-2, pp. 197–216. Masaferro, J. L., Bulnes, M., Poblet, J., and Eberli, G. P., 2002, Episodic folding inferred from syntectonic carbonate sedimentation: the Santaren anticline, Bahamas foreland. Sedimentary Geology, v. 146, pp. 11–24. McHone, J. G., 2000, Non-plume magmatism and rifting during the opening of the Central Atlantic Ocean. Tectonophysics, v. 316, pp. 287–296. McKenzie, D. P., 1978, Some remarks on the development of sedimentary basins. Earth and Planetary Science Letters, v. 40, pp. 25–32. McMillan, N. J., 1973, Shelves of Labrador Sea and Baffin Bay, Canada, in McCrossan, R.G. ed., The future petroleum provinces of Canada; Their Geology and Potential, Canadian Society of Petroleum, Calgary, AB (Memoir 1), pp. 473–517. Miall, A. D., Balkwill, H. R., and Hopkins, W. 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CHAPTER 15

Depositional Evolution of the Gulf of Mexico Sedimentary Basin William E. Galloway

Contents 1. Introduction 2. Crustal Structure and Basin Origin 2.1. Subsidence mechanisms and history 3. Structural Framework 3.1. Basement structures 3.2. Gravity tectonic structures 3.3. Growth structure domains 3.4. Structural growth history 4. Depositional Framework 4.1. Depositional episodes and sequences 5. Depositional History and Paleogeography 5.1. Middle Jurassic–Earliest Cretaceous (Bathonian–Berriasian) depositional episodes 5.2. Early Cretaceous (Valanginian–Cenomanian) depositional episodes 5.3. Late Cretaceous (Cenomanian–Maastrichtian) depositional episodes 5.4. Cenozoic depositional episodes 5.5. Laramide depositional episodes 5.6. Middle Cenozoic volcanism and related depositional episodes 5.7. Miocene depositional episodes 5.8. Early Pliocene–Quaternary depositional episodes 6. Patterns and Generalizations in Gulf Depositional History 6.1. Sediment supply: Sources and drainage history 6.2. Climate and oceanography 6.3. Continental margin evolution 7. Energy Resources Acknowledgments References

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Abstract The Gulf of Mexico is a small ocean basin lying between the North American plate and the Yucatan block. Following initiation in the Middle Jurassic, sea-floor spreading continued approximately 25 Myr. Spreading was asymmetric, creating a broad area of attenuated transitional continental crust beneath the northern basin. Initially, widespread, thick salt deposits accumulated across much of the basin; mobilization of this salt by subsequent sedimentary loading has created a complex suite of gravity tectonic structures. Most salt is now allochthonous, forming extensive stocks and canopies. By the end of the Mesozoic, thermal subsidence had created a deep basin floor, flanked by continental shelves. The resultant basin contains a succession of Late Jurassic through Holocene strata that is as much as 20 km thick. Sediment supply from the North American continent has filled nearly one-half of the basin since its inception, primarily by offlap of the northern and northwestern margins. Depositional history can be generalized in seven phases: (1) Middle-Late Jurassic evaporite and carbonate deposition in a broad, shallow, restricted to open marine basin. (2) Latest Jurassic-Early Cretaceous sandrich clastic progradation from the northern margins. (3) Late-Early Cretaceous development of a rimmed carbonate shelf. (4) Late Cretaceous mixed clastic and carbonate aggradation of the continental margins. (5) Resurgent Paleogene clastic progradation and filling centered in the NW basin. (6) Miocene progradation and basin filling centered in the central and NE Gulf. (7) Late Neogene climatically and eustatically influenced progradation along the central Gulf margin. In contrast Sedimentary Basins of the World, Volume 5 ISSN 1874-5997, DOI 10.1016/S1874-5997(08)00015-4

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to the broad, progradational sediment wedge of the northern Gulf, the Florida margin is a primarily aggradational carbonate platform.

1. Introduction

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The Gulf of Mexico is a small ocean basin lying between the North American plate and the Yucatan block. It contains within its depocenter a succession of Jurassic through Holocene strata that is as much as 20 km thick. Sediment supply from the North American continent has filled nearly one-half of the basin since its inception, primarily by offlap of the northern and northwestern margins. This chapter will focus on the history of this northern fill. The fundamental geologic principal that ‘‘the present is the key to the past’’ has found wide application and success in the Gulf basin. The modern basin (Figure 1) has a central abyssal plain that generally lies at W3 km depth (Bryant et al., 1991). The eastern Gulf floor is dominated by the morphology of the Late Quaternary Mississippi fan. The continental slope of the northern Gulf margin displays a bathymetrically complex morphology that terminates abruptly in the Sigsbee escarpment to the west and merges into the Mississippi fan to the east (Steffens et al., 2003). The hallmark of the central Gulf continental slope is the presence of numerous closed to partially closed, equi-dimensional, slope minibasins. In contrast, the Florida platform forms a broad ramp and terrace that terminates at depth into the nearly vertical Florida escarpment. The western Gulf margin displays intermediate width and it too is bathymetrically complex. Here, numerous contour-parallel ridges and swales dominate the mid- to lower slope morphology. The modern shelf edge, as reflected by a well-defined increase in basinward gradient, generally lies at a depth of 100–120 m. Landward, the northwestern, northern, and eastern Gulf of Mexico is bounded by broad, low-gradient shelves that range from 100 to 300 km in width (Figure 1). Today, and throughout its history, the Florida and Yucatan platforms, which bound the basin on the east and south, persist as sites of carbonate deposition. On shore, the northern and northwestern Gulf margin displays a broad coastal plain (Figure 1). The lower coastal plain, a flat, low-relief surface, is underlain by Neogene and Quaternary strata. The upper coastal plain displays modest relief of less than about 100 m created by Quaternary incision into older Neogene, Paleogene,

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Figure 1 Principal physiographic elements of the Gulf of Mexico basin and adjacent North America.White outline shows approximate geological limits of the Gulf basin. Bathymetry from GEBCO (2003,The GEBCO digital atlas-Centenary edition, British Oceanographic Data Centre); topography from International Centre for Tropical Agriculture (CIAT) (2005),Void-¢lled seamless SRTM dataV2, http://srtm.csi.cgiar.org). Image created using IVS3D Fledermaus.

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and Late Cretaceous strata by numerous large and small rivers. The basin is bounded by a variety of Cenozoic, Mesozoic, and remnant Paleozoic uplands, including the Sierra Madre Oriental of Mexico, the Trans-Pecos mountains of West Texas, the Lower Cretaceous limestone-capped Edwards Plateau, Ouachita Mountains of southern Arkansas (Miall, Chapter 8, this volume), and the Cumberland Plateau and southern Appalachian Mountains of northern Mississippi and Alabama (Ettensohn, Chapter 4, this volume). The northeast Gulf basin merges into the southern Atlantic coastal plain across northern Florida (Miall et al., Chapter 14, this volume); however, the structural basin boundary is generally placed near the current west coast of the Florida peninsula. The geology of the Gulf of Mexico has been reviewed by numerous authors. Two syntheses stand out. Grover Murray’s 1961 Geology of the Atlantic and Gulf Coastal Province of North America summarized classic mid-20th century stratigraphic and structural understanding of the basin. The Geological Society of America’s 1991 Geology of North America volume J, The Gulf of Mexico Basin, edited by Amos Salvador, provided a synthesis of all facets of basin geology and resources integrated through the initial applications of modern concepts of crustal tectonics, depositional systems, genetic stratigraphy, deep-marine studies, and gravity tectonics. The objective of this chapter is to incorporate the wealth of new ideas and information that has been published in the decade since the GSA volume into a succinct description of the stratigraphic framework and depositional history of the northern margin and related deep Gulf of Mexico basin.

2. Crustal Structure and Basin Origin The Gulf of Mexico basin was created by an episode of crustal extension and sea-floor spreading during the Mesozoic breakup of Pangea (Salvador, 1987; Sawyer et al., 1991; Buffler and Thomas, 1994; Harry and Londono, 2004; Jacques and Clegg, 2002). Origin of the basin is reflected in the distribution and nature of the basement crust (Figure 2). The Gulf basin is largely surrounded by normal continental crust of the North American plate. Most of the structural basin is underlain by transitional crust that consists of continental crust that was stretched and attenuated by Middle to Late Jurassic rifting. Two types of transitional crust are differentiated (Figure 2). The basin margin is underlain by a broad zone of thick transitional crust, which displays modest thinning and typically lies at depths between 2 and 12 km subsea depth (Sawyer et al., 1991). The area of thick transitional crust consists of blocks of near-normal thickness continental crust separated by areas of stretched crust that has subsided more deeply. The result is a chain of named arches and intervening embayments and salt basins around the northern periphery of the Gulf basin (Figure 2). Much of the present inner coastal plain, shelf, and continental slope is underlain by relatively homogeneous thin transitional crust, which is generally less than half of the 35 km thickness typical of continental crust and is buried to depths of 10–16 km below sea level. More recent reconstructions of deep seismic traverses (Peel et al., 1995) indicate that basement may lie below 20 km in the central depocenter beneath the south Louisiana coastal plain and adjacent continental shelf. The deep, central Gulf floor is underlain by an arcuate belt of basaltic oceanic crust that was intruded during Late Jurassic through Early Cretaceous sea-floor spreading. The exact nature and actual distribution of this crust is problematic; Figure 2 illustrates the general shape and distribution of oceanic crust suggested by most authors. That central Gulf crust lacks the magnetic signature typical of oceanic crust, compounds interpretation difficulties. The broad history of plate tectonic movements that culminated in the Gulf basin is generally understood (Marton and Buffler, 1999; Pindell and Kennan, 2001; Jacques and Clegg, 2002; Harry and Londono, 2004; Bird et al., 2005), if not fully agreed upon in detail. The Gulf of Mexico opened by the separation of the North and South American plates as rifting spread southward along the Atlantic spreading ridge. A long period of Late Triassic through Early Jurassic extension that created a series of basement grabens and half grabens filled with terrestrial red beds and volcanics presaged the main phase of Late Jurassic–Early Cretaceous Gulf rifting. Recognition of potential seaward-dipping reflectors in the northeastern Gulf suggests an early phase of subaerial volcanism during the initial spreading phase (Imbert, 2005). Continued stretching in Bathonian and Callovian time initiated a broad sag, which opened to the Pacific Ocean. Widespread deposition of thick Louann Salt and associated evaporites, a defining event for the later structural evolution of the Gulf sedimentary fill, spread across the shallow, hypersaline basin centered above the thinned continental crust. Salt thickness was greatest above the marginal crustal embayments and basins and, regionally, above the evolving thin transitional crust (Figure 2). The regional unconformity beneath the evaporite layer separates localized syn-rift from blanket post-rift deposits and is widely taken as the base of the Gulf of Mexico sedimentary basin fill (Sawyer et al., 1991; Buffler and Thomas, 1994). Rotational spreading along a generally east-west trend extending from a pole centered beneath NW Cuba continued through Late Jurassic to as late as the Valanginian (Pindell and Kennan, 2001). Opening of the Gulf entailed approximately 500 km of extension accompanied by southward migration and counter-clockwise

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Figure 2 Crustal types, depth to basement (km), and original distribution of Jurassic Louann pre-marine evaporite beneath the Gulf of Mexico basin. Principal basement structures include: SrA, Sarasota arch; TE,Tampa embayment; MGA, Middle Ground arch; AE, Apalachicola embayment; WA,Wiggins arch; MSB, Mississippi salt basin; MU, Monroe uplift; NLSB, North Louisiana salt basin; SA, Sabine arch; ETB, East Texas basin; SMA, San Marcos arch; RGE, Rio Grande embayment; TA,Tamaulipas arch. Modi¢ed from Sawyer et al. (1991). Note that modern reconstructions suggests crustal depths of W20 km beneath the northcentral Gulf depocenter.

rotation of the rigid Yucatan block to its present position and by extensive NNW-SSE shear along the west flank of the basin (Marton and Buffler, 1999; Pindell and Kennan, 2001; Jacques and Clegg, 2002). Crustal rupture and emplacement of basaltic crust began by the Oxfordian and continued until the termination of spreading in the latest Berriasian or Early Valanginian. Salt deposition ended with onset of sea-floor spreading, and the Louann salt basin was split into northern and southern Gulf segments overlying transitional crust (Figure 2). Jacques and Clegg (2002) suggest two phases of rotation about differing poles. With shift of further inter-plate spreading to the Atlantic and proto-Caribbean basins, cooling and subsidence of the stretched continental and oceanic crustdominated basin development. By the end of the Early Cretaceous, combined deposition of rimming carbonate platforms and subsidence had created the modern outline and morphology of the Gulf Basin (Winker and Buffler, 1988). Late Cretaceous and, especially, Cenozoic history was dominated by loading subsidence, complicated by intrabasinal gravity tectonics. The history of Gulf spreading created four distinctly different basin margin types. The northern margin is a relatively simple divergent margin with a broad zone of stretched continental crust separating oceanic and continental crust. The Yucatan margin, to the south, is also a divergent margin, but juxtaposes thick transitional crust closely to the oceanic crust. This pronounced asymmetry suggests a simple-shear model for extension

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(Marton and Buffler, 1993; Watkins et al., 1995). The Mexico and Florida margins primarily reflect displacement of crustal blocks along a series of transfer faults. To the west, the crustal boundary is characterized by an elongate gravity high and narrow zone of primarily Late Cenozoic growth faults (Ambrose et al., 2005; Bird et al., 2005). On the east, the margin was formed by a series of rhombohedral crustal blocks that rotated between two parallel transfer faults that generally conform to the Florida-Bahamas and the Cuban Fracture Zones (MacRae and Watkins, 1996). The family of basement arches and sags that extends from the Mississippi Salt Basin southeast to the Sarasota Arch (Figure 2) were produced in this transtensional domain (Watkins et al., 1995; Marton and Buffler, 1999; Pindell and Kennan, 2001; Stephens, 2001).

2.1. Subsidence mechanisms and history Like other oceanic basins, total subsidence of the Gulf basin is the sum of crustal stretching, cooling, and loading subsidence. Combined stretching and cooling as crust migrated away from the axial spreading center and then cooling after spreading ceased caused a total tectonic subsidence of 5–7 km of the central thin transitional and oceanic crust (Sawyer et al., 1991). Initially, stretching and cooling subsidence created a starved basin that subsided more rapidly than sediment was supplied. The marine basin expanded and deepened. Subsequent depositional loading of the crust, which soon followed and has continued through the Holocene, has further depressed the crust to its current 10–20 km (Figure 2) below sea level. Loading subsidence has dominated Cenozoic history of the basin. Additional Mesozoic and Cenozoic tectonic phases have further influenced local to sub-regional subsidence history of the Gulf. Several of the marginal highs, including the San Marcos arch, Sabine arch and Monroe uplift display short pulses of uplift of as much as a few hundred meters, creating angular unconformities in Middle Cretaceous and Lower Eocene strata (Laubach and Jackson, 1990). These pulses generally correlate to phases of Laramide thrusting, in turn related to changing rates of Pacific margin plate convergence and changing intracratonic compressional stress. Extensive crustal heating across northern Mexico and the southwestern United States (Gray et al., 2001) uplifted and tilted Mesozoic and Early Cenozoic strata of the western Gulf. Cenozoic mobilization of thick bodies of intrabasinal salt has created as much as 1–2 km of often rapid subsidence of the overlying outer shelf and upper slope sediments at numerous times along segments of the northern Gulf continental margin (Diegel et al., 1995; Galloway et al., 2000). Such salt evacuation has been a major process for creation of local to regional sediment accommodation volume.

3. Structural Framework The depositional history of the Gulf of Mexico Basin is best understood in the context of both the basement structure, which subtly influenced sediment supply and accumulation patterns, and gravity tectonic structure, which reflects dynamic interactions among depositional loading, sediment and salt mobilization, creation or loss of accommodation space, and deformation.

3.1. Basement structures Basement structures and their influence on overlying stratigraphy are most readily apparent around the periphery of the basin underlain by thick transitional crust. They include the halo of embayments (epicratonic basins that open to the central Gulf ) and basins and intervening arches and uplifts (Ewing, 1991) (Figure 2). The basins and embayments typically contain a significant thickness of Louann salt and thicker sequences of Jurassic and Early Cretaceous strata relative to the adjacent arches and uplifts. Salt-floored basins, including the East Texas basin, North Louisiana salt basin, Mississippi salt basin, and Appalachacola embayment (also known as the DeSoto Canyon salt basin) contain well-described families of salt domes and related structures (e.g., Seni and Jackson, 1984). Deep crustal structures of the thin transitional and oceanic crustal domains are less easily defined. Gravity and magnetic data, changes in basement topography and rates of subsidence, and salt distribution all suggest a family of NW-SE trending basement transfer faults created during Atlantic and Gulf extension and spreading phases (Watkins et al., 1995; Huh et al., 1996; Stephens, 2001). In the Late Cretaceous (60–100 Ma), intrusive and extrusive volcanism occurred around the northern and northwestern periphery of the Gulf Basin (Byerly, 1991; Stephens, 2001). Principal volcanic clusters lie around the inner edge of the central and south Texas coastal plain, and in southern Arkansas and the adjacent Monroe

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uplift of northern Louisiana and adjacent Mississippi. Igneous lithologies include basalt, nephelene syenite, phonolite, and peridotite.

3.2. Gravity tectonic structures The Gulf of Mexico basin fill displays one of the best-described and most complex assemblages of gravity tectonic structures to be found in the world (Worrall and Snelson, 1989; Nelson, 1991; Diegel et al., 1995; Jackson, 1995; Peel et al., 1995; Watkins et al., 1996a; Jackson et al., 2003). The combination of a thick, basin-flooring Louann salt substrate, rapid sediment loading, and offlap of a high-relief, continental-margin sediment prism has resulted in mass transfer of salt and overpressured mud up section and basinward throughout Gulf history. The resultant panoply of structures and related features includes: 1. Growth-fault families and related structures (Winker, 1982; Watkins et al., 1996b). Growth faults tend to nucleate and grow during active deposition at the continental margin. Here, extension results from basinward gravitational gliding or translation of the sediment wedge along a detachment zone, typically found within salt or overpressured deep-marine mud (Rowan et al., 2005). Extension creates a family of features, including primary synthetic growth faults, splay faults, antithetic faults, and rollover anticlines (Figure 3A). 2. Allochthonous salt bodies, including salt canopies and salt sheets (Diegel et al., 1995; Fletcher et al., 1995; Peel et al., 1995; Jackson et al., 2003). Loading of the Louann salt has resulted in regional extrusion of salt basinward and up section. Salt canopies typically develop beneath the continental slope, where salt rises as a series of coalescing diapirs or as injected tongues. Salt may also be extruded to the surface, forming salt sheets, or nappes, that move basinward much like salt glaciers. 3. Salt welds ( Jackson and Cramez, 1989; Jackson et al., 1994). Welds (Figure 3B and C) are surfaces that juxtapose discordant stratigraphies. They form where nearly complete expulsion of salt stock feeder dikes, salt tongues, or salt canopies has occurred. 4. Roho fault families (Rowan, 1995; Schuster, 1995; Jackson et al., 2003). Lateral salt tongue extension by gravity spreading creates a linked assemblage of extensional faults and compensating, down-slope compressional toe faults, anticlines, and salt injections in the overlying sedimentary cover (Figure 3B). 5. Salt diapirs and their related withdrawal synclines and minibasins (Seni and Jackson, 1984; Rowan, 1995; Fletcher et al., 1995; Rowan and Weimer, 1998; Jackson et al., 2003). In the Gulf-margin basins and embayments, salt diapirs rise directly from the autochthonous Louann ‘‘mother’’ salt. Basinward, depositional loading of salt canopies and sheets beneath shelf and slope areas causes renewed salt stock evacuation, creating EXTENSION Synthetic Splay Fault Faults Rollover

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Figure 3 Typical intrabasinal gravity tectonic structural styles and features of the northern Gulf margin. (A) Linked salt- and shale-based detachments. (B) Salt-based detachment fault system, or Roho structure. (C) Salt-withdrawal minibasin. Modi¢ed from Karlo and Shoup (1998).

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high-relief salt diapirs and intervening depressions (Figure 3C). Progressive salt evacuation creates shifting, localized sites of extreme subsidence and sediment accumulation. Resulting features include (Figure 3C) withdrawal synclines created by local evacuation of salt from diapir flanks, bathymetric depressions, called minibasins, that form local depocenters, turtle structures, and local fault families including down-to-basin ramp faults, counter-regional flap faults, and crestal faults above salt bodies. 6. Basin-floor compressional fold belts (Weimer and Buffler, 1992; Fiduk et al., 1995; Trudgill et al., 1999; Hall et al., 1998). Basinward gravity spreading or gliding along a detachment zone, and resultant updip extension, requires compensatory compression at the toe of the displaced sediment body. Compressional features include anticlinal toe folds and reverse faults (Figure 3A). They commonly form at the base of the slope, but also can extend onto the basin plain where a stepped discontinuity or termination of the decollement layer occurs.

3.3. Growth structure domains The most complex and complete array of gravity tectonic structures lies within the Cenozoic sedimentary wedge of the northern Gulf of Mexico basin (Figure 4). Principal structural features include an inboard series of strikealigned growth-fault families beneath the coastal plain, complex fault families beneath South Louisiana and its adjacent continental shelf, a broad zone of relatively shallow salt stocks and coalesced autochthonous canopies beneath the continental slope, a base-of slope salt nappe, forming the Sigsbee escarpment, and several sub-slope and basin floor compressional fold belts (Figure 4). Sediment loading of the salt canopy has created a series of largely filled shelf minibasins and closed bathymetric lows, called slope minibasins, on the continental slope. This mosaic of gravity tectonic features can be grouped into genetically related structural domains (Peel et al., 1995) (Figure 5). Each domain had a finite time span of primary growth that can be associated with one or more successive episodes of clastic sediment accumulation in the Gulf. Domains generally become younger basinward, beginning with the Middle Cretaceous Louann detachment domain and culminating in the Plio-Pleistocene minibasin and salt canopy domains of the continental slope. The Oligocene-Lower Miocene and Miocene compressional domains are exceptions to this general pattern. In addition, the full array of gravity tectonic structure domains of the northern Gulf basin includes the salt diapirs and related structures of the East Texas, North Louisiana, Mississippi, and DeSoto Canyon salt basins, which lie around the northern basin periphery, and a series of peripheral grabens, including the Luling–Mexia–Talco, State Line, and Pickins–Gilberton fault zones, that delimit the landward extent of

Figure 4 Structural features of the northern Gulf of Mexico. Compiled fromWatkins et al. (1995) and numerous additional sources.

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autochthonous Louann salt (Figure 5). Growth of structures within these inboard domains occurred largely in Mesozoic time. The three-dimensional structural and stratigraphic architectures of the northern basin are illustrated by a regional N-S section across the north-central basin fill (Figure 6). The boundary between thick and thin transitional crust is reflected by a subsidence hinge that became the focus for development and stabilization of the Cretaceous continental shelf margin, most clearly marked by an extensive reef system. Basinward, the thick Cenozoic sedimentary prism overlies thin transitional crust, which has been depressed more than 16–20 km by sedimentary loading. The prism extends beneath the coastal plain and shelf, reaching its thickest point near the present continental margin. The continental slope extends basinward to about the position of the transitional/ oceanic crust boundary. Beneath this sediment prism, most of the autochthonous Louann salt has been expelled, forming a primary salt weld on the basal Jurassic unconformity that is a principal decollement zone for growth faults. Paleogene and Neogene deposits form an off-stepping series of sediment wedges. Paleocene through Miocene wedges are expanded and deformed by a succession of growth-fault families included within the Wilcox and mixed Upper Eocene and top salt detachment provinces. Cretaceous and Early Tertiary fault extension was accommodated by detachment at the Louann Salt; Oligocene — Recent extension typically detached on allochthonous salt canopies or in marine shales (Rowan et al., 2005). The off-stepping deposition acted as a giant rolling pin, pushing salt upward and basinward into three major salt canopies. The inboard canopy was loaded and largely evacuated by subsequent deposition, forming the vast central Gulf shelf minibasin and roho domains. Beneath the continental slope, a shallow salt canopy forms the slope minibasin and salt canopy domains, which terminate in the Sigsbee scarp. However, at the east end of the slope mini-basin province, salt rose directly from the autochthonous level. The base of the canopy rises through flat-lying basinal Cretaceous and Cenozoic strata to the final salt sheet, which is intruded into Pleistocene strata (Figure 6A). Transects through the NE and NW Gulf margins (Figure 7) illustrate features of additional structural domains and general basin stratigraphy. In the NE Gulf (Figure 7A), the total basin fill is relatively thin, depressing the crust only to depths between 7 and 11 km. The crustal boundary again pins the location of the Mesozoic shelf margin,

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Figure 6 North-south (dip) cross-section of the northern Gulf of Mexico continental margin. (A) Crustal types, generalized stratigraphy, and structural elements including major salt canopies and detachment zones. (B) Principal facies associations (J, Jurassic; K, undi¡erentiated basinal Cretaceous; LK, Lower Cretaceous; UK, Upper Cretaceous; P-E, Paleocene--Eocene; O, Oligocene; M, Miocene; Plio., Pliocene; Pleist., Pleistocene). For location see Figure 5. Modi¢ed from Peel et al. (1995).

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which has been built further basinward only about 50 km by Neogene deposition. Growth faults are few. Limited salt stocks, which rise from the largely evacuated autochthonous Louann, define the eastern margin of the slope minibasin domain. However, the basinal toe of the section illustrates compressional features of the east end of the Miocene compression domain. The NW Gulf transect (Figure 7B) illustrates the structure of the basin depocenter located beneath the continental shelf and slope. The Oligocene–Miocene detachment province is rooted in a decollement located within deep basinal muds of indeterminate age. Basal Louann salt has been evacuated both upward as isolated stocks and basinward to the toe of the continental slope and beyond, leaving a primary weld. In contrast to the central and NE Gulf, the NW Gulf displays broad, complex Middle Cenozoic compressional domains, including the Perdido and Port Isabel fold belts. The Port Isabel fold belt is linked by a decollement to the Miocene Clemente-Thomas, Corsair, and Wanda fault zones of the Oligocene–Miocene detachment province (Figures 4 and 5) (Hall et al., 1998). Like the Mississippi fan fold belt, the Perdido fold belt is located at the depositional limit of basal Louann salt (Fiduk et al., 1995). Additional contraction was accommodated by the compound salt canopy that has been injected up into Oligocene and Miocene section.

3.4. Structural growth history Backstripping of regional cross-sections (Figure 8) reveals the dynamic interplay between deposition, wholesale mass transfer of salt, development of growth structures, and outbuilding of the Gulf margin that has characterized the basin’s history (Diegel et al., 1995; Peel et al., 1995; McBride, 1998). Late Jurassic accumulation of up to 4 km of Louann salt extended across the subsided thinned transitional crust. By the end of the Cretaceous, deposition had loaded and expelled much of the landward part of the autochthonous salt basinward, beneath the paleocontinental slope toe and northern basin floor (Figure 8B). Extension of the upper slope was accommodated by compressional deformation at the slope toe. A remnant layer of autochthonous salt provided the decollement horizon for basinward gravity spreading. By the end of the Oligocene (Figure 8C), successive pulses of Paleogene deposition had prograded the continental margin over the Cretaceous slope, deflating the thick salt under-layer by intrusion of salt stock canopy complexes under the advancing continental slope and further inflation of the abyssal salt sheet. The Oligocene Frio growth-fault zone migrated basinward with the prograding continental margin; here decollement occurred within Upper Eocene mud as well as in the deeper salt. The resultant continental slope was a mix of sediment and near-surface salt bodies. Miocene–Pliocene deposition loaded the salt canopies, triggering passive diapirism and further gravity spreading, creating roho fault systems and isolated salt stocks separated by welds (Figure 8D). Thick minibasin fills separate the salt stocks. Loading also initiated extrusion of a salt sheet at the toe of the slope. Pleistocene deposition has filled updip minibasins and built the continental slope onto the distal salt sheet, where incompletely filled minibasins dominate present slope topography (Figure 8E).

4. Depositional Framework The stratigraphic architecture of the northern Gulf of Mexico Basin displays many elements typical of divergent continental margins (Winker, 1982, 1984; Winker and Buffler, 1988). (1) Above a break-up unconformity, initial strata onlapped the subsiding basin margin. (2) Following this onlap phase, sediment supply overcame subsidence, and margin aggradation accompanied by offlap-dominated. A deep, sediment-starved basin center became separated from the marginal coastal plain and shelf by a clearly defined shelf edge and slope. (3) Further deposition created a succession of offlapping stratal units constructing a broad coastal lain and continental shelf. This nearly continuous depositional record, which covers more than 160 Ma of geologic time and continues today, produced a succession of regionally correlative stratigraphic units that are separated by major marine flooding horizons, sediment-starvation surfaces, and erosional unconformities.

4.1. Depositional episodes and sequences Northern Gulf basin stratigraphic framework, chronology, and nomenclature were established during the earlyto mid-20th century using conventional stratigraphic concepts. The thick, monotonous, siliciclastic Cenozoic section was subdivided using the fossiliferous marine shale tongues that record regional transgressions across the northern basin. This concept of transgression-bounded genetic units was formalized in a seminal paper by D.E. Frazier in 1974. Frazier argued that the Gulf Cenozoic fill recorded a succession of depositional episodes, each characterized by a foundation of progradational marine and coastal facies, overlain and replaced landward by

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Figure 8 Reconstruction of the regional north-south cross-section of the Gulf continental margin showing evolution of salt canopies and fault complexes. Modi¢ed from Peel et al. (1995).

aggradational coastal plain and fluvial facies. Both facies successions were capped by a relatively thin succession of transgressive or back-stepping coastal and marine shelf facies. The ‘‘Frazierian’’ genetic unit is bounded basinward by submarine starvation surfaces (condensed beds) created during and soon after transgressive retreat of coastal depositional systems. If relative or eustatic sea-level fall further punctuates the history of a depositional episode, the genetic unit will contain an internal subaerial unconformity within its updip strata. Using the Frazierian depositional model, Galloway (1989a) defined the genetic stratigraphic sequence as a fundamental unit of Gulf of Mexico Cenozoic stratigraphy. The genetic sequence consists of all strata deposited during an episode of sediment influx and depositional offlap of the basin margin. It is bounded by a family of surfaces of marine non-deposition and/or erosion created during transgression, generalized as the maximum flooding surface. This pattern is readily recognized in the Paleogene section, where transgressive marine shelf mudstone and glauconitic sandstone units extend to outcrop (Galloway, 1989b). It also applies in Neogene strata,

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where prominent transgressive markers record glacioeustatic sea-level rise events (Galloway et al., 2000). Thus, genetic sequences typically correspond closely to widely used stratigraphic nomenclature. The depositional sequence paradigm, which uses subaerial erosion surfaces as sequence boundaries, provides an alternative to the traditional Gulf basin lithostratigraphic framework and has been applied by several authors (e.g., Yurewicz et al., 1993; Mancini and Puckett, 1995; Lawless et al., 1997) especially to Late Neogene strata that are strongly influenced by glacioeustasy (Weimer et al., 1998; Roesink et al., 2004). Depositional sequence models for carbonate and mixed successions, which are appropriate for the Mesozoic Gulf fill, are summarized and illustrated by Handford and Loucks (1993). The synthesis of Gulf depositional history and physical stratigraphy as presented here largely utilizes the traditional lithostratigraphic framework of the Mesozoic and Paleogene sections and the regional marine flooding horizons characterized by widely identified faunal markers within Neogene strata. Building upon the syntheses of Winker and Buffler (1988), Galloway (1989b), Morton and Ayers (1992), and Galloway et al. (2000), I propose a genetic stratigraphic framework that groups strata into a succession of 29 principal Gulf of Mexico depositional episodes (Figures 9–12). First and foremost, each episode records a long-term (ca. 2–12 Ma) cycle of sedimentary infilling, typically accompanied by shelf-margin offlap, of the northern Gulf basin. Deposits of each episode are characterized by lithologic composition (sandstone, mudstone, carbonate, evaporite), vertical stacking of lithofacies and parasequences, and relative stability of sediment dispersal systems and consequent paleogeography. Almost all of the depositional episodes terminated with a phase of deepening and/or basin-margin transgression (Figures 10 and 12). Deposits of episodes are bounded by prominent, widely recognized, and well-documented stratigraphic surfaces (Figures 10 and 12). Bounding surfaces variously include marine starvation and condensed horizons, maximum flooding surfaces, marine and subaerial erosional unconformities, and faunal gaps that are described and interpreted by multiple authors. Such depositional episodes conform to the basic definition of a sequence as a contiguous suite of genetically related strata bounded in part by unconformities. In fact, most of the Mesozoic depositional episodes described here correspond to seismic or depositional sequences identified by one or more authors (e.g., Yurewicz et al., 1993; Dobson and Buffler, 1997; Goldhammer and Johnson, 2000). They are widely recognized as fundamental stratigraphic building blocks of the basin fill. At the same time, a depositional episode framework is sufficiently flexible and robust to accommodate stratigraphic units that were variously dominated by tectonic deformation, sediment supply and composition histories, or eustatic sea-level change.

5. Depositional History and Paleogeography The stratigraphy, depositional system framework, and paleogeographic evolution of the northern Gulf basin will be discussed in the context of the 29 depositional episodes. These episodes logically cluster into Bathonian– Berriasian (Middle–Late Jurassic and earliest Cretaceous), Early Cretaceous, Late Cretaceous, and Cenozoic families. Each episode is recorded by a genetic sequence of strata that is constructed of the facies of a suite of carbonate and/or terrigenous clastic depositional systems. These systems, in turn, record geologically long-lived paleogeographic features that constituted the physical geography of the northern Gulf of Mexico. The depositional system classifications (Figure 13) follow those of Galloway and Hobday (1996) and Handford and Loucks (1993).

5.1. Middle Jurassic–Earliest Cretaceous (Bathonian–Berriasian) depositional episodes The Upper Jurassic and lowest Cretaceous Louann, Norphlet, Smackover, and Cotton Valley episodes form a tectonostratigraphic megasequence bounded below by the break-up unconformity and above by a prominent intra-Valanginian unconformity, which records the termination of sea-floor spreading (Todd and Mitchum, 1977; Winker and Buffler, 1988; Wu et al., 1990; Salvador, 1991b; Dobson and Buffler, 1997; Marton and Buffler, 1999). Initial breakup created a shallow Gulf basin with a connection to the Pacific Ocean across central Mexico. Widespread deposition of Louann salt and associated anhydrite blanketed subsiding transitional crust (Salvador, 1987, 1991a, 1991b; Dobson and Buffler, 1997). As much as 4 km of nearly pure halite, deposited over a span of almost 10 Ma, buried the underlying topography and onlapped northward onto the structural margin of the Gulf (Salvador, 1987) (Figure 2). Salt accumulation was replaced, in the Oxfordian, by deposition of a relatively thin, widespread siliciclastic-dominated sequence that is best known around the northern and northwestern Gulf margin as the Norphlet Formation. The boundary between the Louann and Norphlet sequences is poorly

517

Depositional Evolution of the Gulf of Mexico Sedimentary Basin

Time (Ma)

Stages Maastrichtian

Depositional Architecture Escondido

Navarro

Olmos - Nac.

70

Taylor Ariacacho

90

Santonian Coniacian Turonian Cenomanian

Eagle Ford/U. Tuscaloosa Woodbine / Tuscaloosa Buda Kiamichi

Washita Georgetown

CRETACEOUS

110

Austin / Eutaw

Albian

Glen Rose

Stuart City

Edwards Paluxy

Fredbg. F.L.

G.R. Bexar James

Pearsall Aptian

Pine Island

Early

120

Sligo

Sligo

Barremian Hauterivian

130

RIMMED SHELF

Late

80

100

OPEN SHELF

San Miguel Campanian

Hosston

Valanginian Knowles

140

Cotton Valley Tithonian

170

JURASSIC

Kimmeridgian Oxfordian Callovian

Middle

160

Late

150

Formation of Oceanic Crust

Bossier Gilmer Smackover

Buckner

Basalt RAMP

Berriasian

Norphlet Louann Salt

Bathonian Bajocian

Connection opened to Western Interior Seaway

Figure 9 Generalized Mesozoic stratigraphic succession and architecture of the Northern Gulf of Mexico basin. Time scale of Berggren et al. (1995). Modi¢ed fromWinker and Bu¥er (1988).

defined; deposition may have been continuous or disconformable (Salvador, 1991a). In either case, the Norphlet deposits further onlapped the break-up unconformity, especially in the structural embayments of the northeast Gulf margin. There, several small alluvial fan, braidplain, and delta systems created local depocenters up to 300 m thick. Eolian, sabkha, and playa deposits are also abundant, indicating continued aridity. Basinward, siliciclastics grade into marine shale and limestone. Although I have differentiated the Louann and Norphlet as two episodes, based on the prominent lithologic change and evidence of a pulse of clastic input, the Norphlet might alternatively be considered the transgressive cap of a single, evaporite-dominated Louann sequence (Goldhammer and Johnson, 2000). Continued Oxfordian transgression onto the stable basin margin initiated the first carbonate-dominated depositional episode of the Gulf. Together, the Smackover, Buckner, and Gilmer Formations record a ca. 5 Ma cycle generally bounded above and below by transgressive flooding surfaces (Salvador, 1991b; Prather, 1992; Dobson and Buffler, 1997; Goldhammer and Johnson, 2000; Mancini and Puckett, 2005) (Figure 9). Initial

518

William E. Galloway

Time (Ma)

Stages

Depositional Episodes Clastic Supply Carbonate Platform

Major Surfaces D

Maastrichtian Nacatoch Olmos

70

Composite Episodes Navarro

MFS

San Miguel

U. Taylor

Campanian

D

LATE

80

L. Taylor D Santonian Coniacian

90 Turonian Cenomanian

110

MFS

Tusc.-

Albian

Woodbine

Tuscaloosa Woodbine Washita

D Fredericksburg

Paluxy

D Glen Rose

EARLY

Tuscaloosa Eagleford

D

L. Stuart City

James 120

Austin

D D

U. Stuart City

CRETACEOUS

100

D

Eutaw

D

MFS

Aptian

Glen Rose

D MFS

James

D

Sligo

Sligo

Barremian Hosston

MFS

Hauterivian

130

D Lower Hosston

Valanginian 140

Tithonian

JURASSIC

D Cotton Valley

Cotton Valley D

Kimmeridgian Oxfordian Callovian

MIDDLE

170

LATE

150

160

Knowles

Berriasian

Haynesville Norphlet

MFS Smackover

Smackover MFS

Norphlet Louann

Bathonian

Bojocian

Figure 10 Mesozoic depositional episodes as re£ected by major phases of siliciclastic and carbonate sediment accumulation in the northern Gulf basin. Major stratigraphic surfaces include basin-margin unconformities, deepening events (D) and associated ravinement, and maximum £ooding disconformities (MFS). Composite episodes re£ect regionally concordant stratigraphic units bounded by major surfaces and a relatively stable paleogeography.

deposits consisted of fine-grained, dark, carbonate ramp sediments, which were succeeded by a heterogeneous assemblage of carbonates, including prominent ramp-edge grain shoals. These banks aggraded and coalesced to form a broad shoal system around the northwest and west-central Gulf (Figure 14) (Budd and Loucks, 1981; Moore, 1984). In the northeastern Gulf, grain shoals formed around the emergent basement arches. In mid-episode, evaporites of the Buckner Formation accumulated on the shoal-restricted, shallow inner platform. Seaward, carbonate muds formed a broad carbonate ramp, or, to the east, a nascent carbonate slope. Clastic influx was minor. Small delta and flanking shore-zone systems (the Haynesville Formation) prograded onto the

519

Depositional Evolution of the Gulf of Mexico Sedimentary Basin

Lower Miocene

L

OFFLAP

Upper Miocene Middle Miocene

E

20

Basin Margin Pinch Out

M

10 Miocene

Depositional Architecture Pleistocene Pliocene Bulminella 1 Central Dep.

L

Stage Pleistocene Pliocene

EL

Time (Ma) 0

Queen City

PERCHED

Eocene

Sparta

NW Depocenters

E L

Jackson Yegua

M

40

Frio

Oligocene CENOZOIC

30

Midway

OFFLAP RMP .

E

Lower Wilcox

Paleocene E

60

Upper Wilcox

L

50

Figure 11 Generalized Cenozoic stratigraphic succession and architecture of the Northern Gulf of Mexico basin. Time scale of Gradstein et al. (1995).

northeastern Gulf margin. The episode ended with terminal flooding and deposition of the transgressive Gilmer Limestone. The pulse of clastic sediment input along the northeastern Gulf margin (Figure 10), which coincided with the later part of the episode, limited transgressive Gilmer carbonate deposition to the outer ramp and basin. Sandstones of the Cotton Valley depositional episode (Figures 9 and 10) abruptly overrode the transgressive Gilmer and Haynesville strata (Salvador, 1991b; Prather, 1992; Dobson and Buffler, 1997; Goldhammer and Johnson, 2000; Klein and Chaivre, 2002). Locally, patterns of reflection and stratal terminations suggest the presence of a disconformity associated with transgression or maximum flooding or clastic burial of the Smackover ramp. The dramatic change from carbonate-dominated to siliciclastic-dominated deposition across the entire northern Gulf basin indicates that continental uplift or climate change rejuvenated adjacent North American source areas. Large, sandy delta systems prograded from major fluvial axes centered in the East Texas basin, Mississippi salt basin, and Apalachicola embayment. Suspended sediment spread basinward to form a broad, muddy, marine shelf platform that built basinward beyond its older Jurassic foundations. As deposition progressed, a distinct shelf/slope break emerged. On this platform, marine reworking connected the delta systems with sandy shore-zone and shelf systems. This major episode of clastic input and progradation lasted more than 10 Ma, and deposited more than 300 m of sediment around much of the northern Gulf. It terminated with a relatively brief phase of carbonate accumulation, creating the back-stepping Knowles Limestone (Figure 9). This carbonate blanket marks the terminal transgression of a clastic-dominated episode; together, the Cotton Valley and Knowles form a major transgression-bounded sequence. Although the Cotton Valley depositional episode ended with the conventional record of transgression, its deposits are separated from strata of the overlying Lower Cretaceous Hosston episode by a singularly prominent unconformity throughout the northern Gulf divergent margin (Salvador, 1991b; Goldhammer and Johnson, 2000) (Figure 10). Updip, this unconformity records the entire Valanginian (about 5 Ma); basinward, until strata become concordant beyond the Cotton Valley progradational margin. Here, Valanginian strata form a fore-shelf lowstand wedge (Figure 9). The unconformity records subaerial exposure and erosion, which clearly reflect progressive uplift and basinward tilting of the northern Gulf margin. Coincidence of the unconformity with termination of sea-floor spreading in the Gulf and its medial location within a 25 Ma phase of coarse clastic sedimentary influx to the northern basin indicate that it is a direct consequence of intraplate stress regime changes

520

William E. Galloway

Time (Ma) 0

Stage

Clastic Supply

Major Surfaces

Pleistocene

Pleistocene E L

Pliocene

Plio. Bul. 1

L

Upper Miocene

Middle Miocene

Middle Miocene

10

MFS

E

20

Bul. 1

Upper Miocene

M

MFS

Miocene

Composite Episodes

Lower Miocene

Lower Miocene

L

MFS

Frio

Frio / Vicksburg E

Cenozoic

Oligocene

Jackson

L

30

Jackson MFS Yegua

Yegua MFS

40 M

Sparta

Eocene

Sparta MFS

Queen City

Queen City MFS

E

50 Upper Wilcox

Upper Wilcox

L

MFS

Paleocene

Lower Wilcox

E

60

MFS Lower Wilcox

MFS

Figure 12 Cenozoic depositional episodes as re£ected by major phases of siliciclastic sediment accumulation in the northern Gulf basin. Major stratigraphic surfaces include basin-margin unconformities and maximum £ooding disconformities. Composite episodes re£ect regionally concordant stratigraphic units bounded by major surfaces and a relatively stable paleogeography. Neogene episodes incorporate multiple glacioeustatic cycles and their resultant high-frequency sequences.

and resultant deformation of the North American plate. Together with the sub-salt unconformity, the Valanginian unconformity bounds the syn-drift strata of the early Gulf.

5.2. Early Cretaceous (Valanginian–Cenomanian) depositional episodes Following termination of Gulf spreading, a succession of six composite depositional episodes (Figure 10, Lower Hosston–Washita) provides a record of diminishing continental source area relief and basin-margin stabilization (Winker and Buffler, 1988; McFarlan and Menes, 1991; Scott, 1993; Yurewicz et al., 1993; Marton and Buffler, 1999; Goldhammer and Johnson, 2000; Kerans and Loucks, 2002; Badali’, 2002; Mancini and Puckett, 2005). The climatic setting remained tropical and arid. Clastic input decreased and carbonate deposition came to dominate the northern GOM (Figure 9). Two phases of regional progradation of the reef-rimmed carbonate margin, separated by a regional Early Albian flooding event (Figure 9), produced a well-defined shelf edge separating open to restricted, shallow platform depositional systems from steep slope and deep basinal equivalents. Following this distinctive phase of Early Cretaceous deposition, which lasted for nearly 40 Ma, the intraCenomanian unconformity and subsequent resurgent clastic deposition marked a basin-scale reorganization of regional depositional patterns. Continental uplift and erosion that supplied clastics was focused on the Mississippi embayment and has been associated with subcrustal passage of the Bermuda hotspot (Cox and Van Arsdale, 2002).

521

Depositional Evolution of the Gulf of Mexico Sedimentary Basin

Delta

CO 3 Shor e Zo n e

Zone ore

Sh

Br

Mixed Shelf

aid

Grain Shoal

De

lta

Coastal Plain

pen Shelf

O

lina

Sa

Reef Carbonate Ramp bris De ron p A

Rimmed Shelf

Slo

pe

Abyssal Plain Fan A

B Plain

one

re Z

Sho

Shelf

l

ta as

Co

ne Zo e or S

Sh

R.A.

Ra

elf Sh

mp

Shelf Fed Apron

Delta - Fed Apron

Sc

arp

Fan De lt Ap a Fe ron d

av e De - D lta om . W

an

dy

CO3 Shelf

Fluvial - Dominated Delta

Abyssal Plain Drift

Figure 13 Generalized paleogeographies of (A) carbonate-dominated and (B) siliciclastic-dominated episodes of deposition within the northern Gulf of Mexico. Principal depositional systems are distinguished using this format on the following paleogeographic maps.

Clastic sediment supply continued to dominate Valanginian and Hauterivian deposition. The conglomeratic, sandy Hosston (eastern Gulf margin) and Travis Peak (Texas) formations record this siliciclastic influx. Hosston stratigraphy is complex, however, and displays three depositional styles. Basal Hosston deposits form a shelfmargin prograding wedge that records coastal plain and shelf bypass during formation of the Valanginian unconformity (Yurewicz et al., 1993). Beginning in Late Valanginian, Lower Hosston strata were buried as subsidence of the basin margin and northward expansion of the Gulf basin resumed (Figure 9). During the Hauterivian, Hosston strata onlapped the northern Gulf margin, aggraded the shelf, and prograded the shelf edge basinward of the lowstand wedge. In the Late Hauterivian, deepening and transgressive flooding interrupted depositional offlap. Together, the lowstand wedge and overlying aggradational and progradational Lower Hosston

522

William E. Galloway

100°

90° Haynesville Delta

Smac kove r S ho al

m ste Sy

85°

Ramp

S.A.

W.A.

30° o Sl pe

? ?

0

200 mi

0

200 km

95°

90°

Figure 14 Paleogeography and principal depositional systems of the Upper Jurassic Smackover depositional episode. Depositional elements re£ect lower Smackover facies distribution.

100°

95°

90°

85°

Travis Peak Delta Hosston Delta

30°

? ?

0 0

95°

200 mi

Depocenter

200 km

90°

85°

Figure 15 Paleogeography and principal depositional systems of the Lower Cretaceous Lower Hosston depositional episode.

deposits form a compound depositional episode that was initiated by tectonically forced regression and terminated by transgression and onset of carbonate deposition on the outer shelf (Figure 10). Four deltaic depocenters-dominated Lower Hosston accumulation (Figure 15) (McGowen and Harris, 1984; Dutton, 1987; McFarlan and Menes, 1991). Source areas located to the northeast and northwest coalesced into four sandy bedload fluvial systems that prograded marine-modified braidplains and deltas into the Apalachicola embayment, Mississippi salt basin (Hosston delta), East Texas basin (Travis Peak delta), and Rio Grande embayment. Extensive wave reworking of sandy delta fronts nourished a series of interdeltaic strandplain and

Depositional Evolution of the Gulf of Mexico Sedimentary Basin

523

barrier/lagoon systems and associated shallow sandy shelves. Suspended sediment spread from deltas to form a muddy, prograding outer shelf and slope across the northern Gulf. Only the interdeltaic shelf above the San Marcos arch, in central Texas, was carbonate-dominated. Following Late Hauterivian flooding, a mixed carbonate/clastic depositional episode inaugurated the reefrimmed carbonate margin progradation that is the hallmark of the Lower Cretaceous GOM section (Kauffman and Johnson, 1997). Landward, siliciclastic Upper Hosston deposition continued, but progressively decreased in geographic extent through the Barremian (McGowen and Harris, 1984; Bebout et al., 1981, McFarlan and Menes, 1991). At the same time, carbonate-forming open platform, reef, and grain shoal systems of the Sligo Formation expanded landward across the outer shelf and created a well-organized shelf-margin reef system that stretched from northern Mexico to the southern Florida platform (Bebout, 1977). Coral-like rudistids built these and later Lower Cretaceous reefs and banks. After modest progradation, the barrier reef-rimmed shelf margin stabilized and aggraded along a narrow belt located above the boundary between thin and thick transitional crust. This shelf margin, which persisted until the Early Cenozoic, reflects a hinge line between the two crustal subsidence domains (Sawyer et al., 1991). By Early Aptian, the carbonate environments extended to the depositional limits of the basin, reducing clastic facies to a thin, undifferentiated muddy ‘‘ring’’ around the basin fringe. The shallow northern Gulf shelf formed a broad, open carbonate platform upon which local rudistid banks and grain shoals accumulated (Bebout and Loucks, 1974; Bebout et al., 1981). Shoals, patch reefs, and banks are particularly abundant over residual basement highs, such as the Sabine and San Marcos arches. After nearly 10 Ma of carbonate platform growth and consolidation, abrupt Aptian deepening terminated the Sligo depositional episode (Figure 10). The northern Gulf shelf was blanketed by the thin, widespread Pine Island Shale (Figure 9). Temporary shoaling and rejuvenation of carbonate-forming environments resulted in a brief (1–2 Ma) episode of carbonate platform and margin deposition, forming the James Limestone (Figure 9). A second, more extensive drowning event is recorded by the Bexar Shale and terminated the James depositional episode. The shales, thin sandstones, and limestones of the James depositional episode are characteristically dark and fine-grained. Together, the Pine Island–James–Bexar interval, which form the Pearsall Group of the Texas Gulf margin, constitute a punctuated, retrogradational stratigraphic systems tract that culminated with a basinwide flooding surface. Following this flooding, the rimmed margin that was initially re-established by Albian depositional episodes was displaced far landward of the underlying Sligo shelf edge around much of the northern Gulf (Figure 9). Shelf drowning was followed by slow reestablishment of regional carbonate platform and barrier reef systems — the Stuart City reef — that are defining features of the Mid-Cretaceous Gulf basin. Reef progradation and aggradation reconstructed the Cretaceous shelf edge into a nearly continuous barrier rim extending from Mexico to south Florida. Rudistids continued as principal bank and reef builders, forming the barrier reef as well as platform patch reef and bank complexes (Bebout and Loucks, 1974; Scott, 1990; Kauffman and Johnson, 1997). Corals, encrusting algae, and stromatoporoids contributed to reef construction. Bathymetric contrast between the shallow carbonate platform and the deep central Gulf likely approached its maximum. Winker and Buffler (1988) calculated probable central Gulf water depths above oceanic crust of between 4.2 and 4.7 km, deeper than the modern Gulf abyssal plain. Three deepening events and one episode of clastic sediment input punctuated the ca. 12-Ma growth of the Albian rimmed platform, creating three depositional episodes named for the outcropping Glen Rose, Fredericksburg, and Washita groups that compose them (Figure 9). Basal beds of the Glen Rose depositional episode onlap underlying strata (Yurewicz et al., 1993), suggesting a local disconformity. Strata of the Glen Rose episode are characterized by sandy to argillaceous, oolitic and bioclastic lime mudstone, packstone, and grainstone. Contained within the middle of the Glen Rose limestones are evaporites and dolomites of the Ferry Lake Anhydrite, which accumulated in an internal lowland salina behind the barrier reef. Detailed facies analysis of the Glen Rose in the NW Gulf margin indicates an internal flooding surface that might be used to further subdivide it into upper and lower episodes (Kerans and Loucks, 2002). Terminal deepening followed by an updip unconformity, shoaling, and resurgent clastic influx onto the inner shelf separate the Glen Rose from the overlying Fredericksburg depositional episode. The Fredericksburg genetic sequence consists of three principal lithostratigraphic components. The Paluxy (Texas) and Danzler (Mississippi and Alabama) Formations record depositional progradation of deltas and flanking shore-zone systems onto the inner- to middle shelf of the East Texas basin and northeastern Gulf early in the episode (Figure 16) (Caughey, 1977; McFarlan and Menes, 1991). Shelf limestone and dolomite of the Edwards Group and its equivalents accumulated throughout the episode on the outer shelf and transgressed landward over Paluxy inner-shelf, deltaic, and shore-zone systems late in the episode. As clastic bypass to the slope decreased and carbonate systems dominated shelf margin and slope sedimentation, the declivity and relief of the continental slope increased to angles exceeding 101 (Corso et al., 1989). The resultant high-relief, steeply bounded carbonate margin around the northern Gulf set the stage for later development of the prominent Mid-Cretaceous stratigraphic discontinuity.

524

William E. Galloway

90°

100° Paluxy

85°

Delta

City

Dan

Re

zler

ef

s

Del

tas

ar t

D e v il

ough er Tr Riv

St u

30°

McKnight Salina

0 0

200 mi 200 km

Figure 16 Paleogeography and principal depositional systems of the Lower Cretaceous Fredricksburg depositional episode.

Throughout the episode, low-relief reefs flourished along the shelf margin, forming the widely recognized Stuart City reef (Figure 16) (Bebout and Loucks, 1974; Scott, 1990). Above the San Marcos arch and in the northeast Gulf, the Stuart City aggraded on the foundation of the Sligo shelf-margin reef. The reefal shelf margin prograded across the central Gulf, regaining and aggrading the older Sligo margin, except along the segment lying south of the East Texas basin and Sabine arch where the Stuart City reef remained a few tens of kilometers landward of the underlying Sligo reef. At the Rio Grande embayment, the Stuart City reef axis diverted westward, across the embayment, and then inland around the Maverick Basin (a subsiding intrashelf basin), forming the Devil’s River trough (Winker and Buffler, 1988). Here, the reef sufficiently restricted the shelf to form the extensive McKnight salina by Late Fredericksburg deposition (Figure 16). Exact configuration of the Maverick basin changed through the episode, and paleogeographic reconstructions by various authors commonly differ in detail. Fredericksburg episode deposition terminated with widespread accumulation of dark, calcareous claystone and interbedded lime mudstone of the Kiamichi Formation and its equivalents (Figure 9). The Kiamichi lithologies indicate regional deepening of the northern Gulf shelf. Concomitantly, epeirogenic uplift and tilting of the landward basin margin created a minor unconformity at the base of overlying Washita strata. During the Middle Albian, global sea-level rise and ongoing subsidence and northwestward expansion of the Gulf of Mexico combined to open a connection to the Western Interior seaway. It is appealing to suggest that this connection reorganized or diverted continental drainage systems, greatly reducing or terminating sand supply to the northwestern Gulf and leading to the widespread expansion of clean Edwards Group carbonate deposition onto the rapidly shrinking Paluxy coastal plain within the Fredericksburg episode. Clastics continued to be derived by basin-margin streams, but local source areas were of limited area and low relief in the northwestern Gulf. However, in the northern Gulf, deltaic and shore-zone systems fed by streams arising in the eastern uplands continued to accumulate sand and mud throughout the Fredericksburg episode. The Washita depositional episode bridged the Early to Late Cretaceous boundary; however, its depositional style remained that of the Early Cretaceous. The episode was characterized by climax, aggradational growth of the Stuart City reef. On the northern Gulf platform, widespread accumulation of shallow shelf lime mud, bioclastic sand, marl, and calcareous mud-dominated. Although clastic influx in the northwest Gulf was substantial, it was limited to fine, suspended load that was dispersed widely across the shallow marine shelf. Latest deposits of the episode are dominantly marine to restricted marine shale, again recording partial drowning of the carbonate platform and diminished carbonate formation within deeper embayments. The episode terminated with the formation of one of the major discontinuities in the Mesozoic record of the Gulf, the Mid-Cretaceous unconformity, or MCU, which is widely used as a practical boundary between Early and Late Cretaceous rocks in the basin. The Mid-Cretaceous unconformity records a profound break in the

Depositional Evolution of the Gulf of Mexico Sedimentary Basin

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depositional architecture of the northern Gulf of Mexico (Wu et al., 1990; Buffler, 1991). The broad carbonatedominated shelf was replaced by alluvial, deltaic, and coastal depositional systems. The reefal Lower Cretaceous shelf margin, which had persisted for nearly 14 Ma, was abandoned and regionally overstepped by clastic progradation. Subsequently, Upper Cretaceous strata blanketed the relict shelf edge, subduing its morphology and creating a ramp-like inflection across the buried reef complex. The depocenter shifted from the shelf margin and basement-controlled basins and embayments to the fore-shelf continental slope. Along the Florida and western Gulf continental slopes, scours and channel cuts record onset of active submarine erosion. The MCU has been widely attributed to the Mid-Cenomanian sea-level fall of the Haq chart (Buffler, 1991; Yurewicz et al., 1993). However, several attributes of the MCU indicate that global sea-level change was at best a minor factor in its formation. (1) Uplift and tilting of the San Marcos arch, Sabine uplift, and Monroe uplift removed much of the Lower Cretaceous section. Subaerial erosion cut as deeply as the Upper Jurassic Cotton Valley sandstones in northeast Louisiana, indicating uplift of as much as several hundred meters. Angular discordance across the MCU clearly demonstrates the role of tectonic uplift in its origin. Changes in crustal stress regime, likely associated with changing rates of Pacific and North American plate convergence and the Sevier orogeny of the western United States may explain the basin flank deformation (Laubach and Jackson, 1990; Cao et al., 1993). (2) Uplift was coincident with and closely followed by a nearly 8 Ma influx of sandy sediments from fluvial systems draining eastern continental uplands. Clearly uplift and erosion rejuvenated or created new upland sources. Timing and location of post-MCU clastic depocenters is consistent with interpreted uplift of the eastern interior, beneath what is now the Mississippi embayment, which parallels the modern Mississippi Valley, due to passage of the Bermuda hotspot beneath thinned Paleozoic crust (Cox and Van Arsdale, 2002). (3) The MCU can be traced down the bounding continental slopes where it records a variable period of sediment starvation and separates Early Cretaceous basinal carbonates from Late Cretaceous or Cenozoic basinal mudstone (Buffler, 1991). However, its interpreted correlation as a prominent reflection horizon beneath the Gulf floor (Buffler, 1991) has been disproven by recent deep-water drilling (Dohmen, 2002). (4) The stratigraphic context shows that the shallow water carbonate factory was progressively drowned by the shelf deepening, as recorded in the uppermost Washita episode, then poisoned as clastics from rejuvenated fluvial systems poured onto the northern shelf. The MCU is thus an excellent example of a drowning unconformity (Schlager and Camber, 1986; Wu et al., 1990). As regressive clastic systems prograded over the dying Stuart City reef, sedimentary bypass and slumping created onlap relations between the clastic and carbonate slope wedges. Unlike a short-term sea-level fall, differential tectonic uplift of the basin margin, creation of a new upland source area, and tilting subsidence of the outer shelf and shelf margin readily explain concomitant subaerial erosion, long-term rejuvenation of clastic influx, carbonate suppression, and a permanent change in basin-wide depositional style across a composite unconformity surface.

5.3. Late Cretaceous (Cenomanian–Maastrichtian) depositional episodes The Late Cretaceous, above the MCU, contains at least six depositional episodes (Figure 10) (Winker and Buffler, 1988; Wu et al., 1990; Sohl et al., 1991; Mancini and Puckett, 1995; Goldhammer and Johnson, 2000; Liu, 2004). Additional known deepening or transgressive events might be used to further subdivide the section. However, six bounding transgressions associated with regional flooding surfaces or basin-margin disconformities differentiate six clastic supply episodes. The Tuscaloosa/Woodbine composite depositional episode consists of the Lower and Upper Tuscaloosa episodes of the Louisiana margin and the Woodbine and Eagle Ford episodes of the Texas margin. It records major progradational deltaic systems that built along the Mississippi embayment and into the East Texas basin (Figure 17). Ongoing uplift of the Sabine arch separated the two clastic depocenters and dispersal systems. The larger of the two fluvial/deltaic axes deposited the Tuscaloosa Formation. The Tuscaloosa fluvial/deltaic system was rapidly forced across the shelf and spilled over the abandoned Stuart City reef, to create a prograding shelfmargin wedge of delta and delta-fed slope apron sandstone and mudstone (Mancini et al., 1987). The prograding deltas constructed a new shelf edge slightly seaward of the foundered reef. Offlap of the clastic wedge, which was more than 1-km thick, onto the steep carbonate slope initiated the first of the many growth-fault families of the northern Gulf (Figure 4). Tuscaloosa regression was interrupted by transgression, creating lower fluvial/deltaic and upper deltaic sandstone units separated by the ‘‘marine Tuscaloosa’’. Through later Cenomanian, Tuscaloosa deltas backstepped as the thermal uplift began to collapse and sediment supply decreased. To the west, the Woodbine fluvial/deltaic system remained largely on the shelf (Figure 17). The delta, which was wave-dominated, prograded to the southwest into the East Texas basin (Oliver, 1971; Turner and Conger, 1984). However, distal suspended mud spread across the Cretaceous reef and built a muddy shelf margin that merges with the Tuscaloosa deltaic wedge south of the Sabine arch. The Woodbine sediment dispersal system records a single clastic pulse that was, however, complicated by ongoing uplift and subaerial exposure of the

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Sabine arch. Emergence of the uplift first provided a local source area and then culminated in angular truncation of Woodbine strata on the east flank of the East Texas basin, creating the unconformity trap for the giant East Texas oil field. Transgression of the Woodbine fluvio-deltaic plain, which lay more distant from the upland source and inferred hotspot-created uplift, relatively early led to widespread deposition of the Eagle Ford Shale across a broad muddy shelf that was contemporaneous with renewed progradation of the Upper Tuscaloosa (Figure 10). The northwest Gulf remained fully open to the Cretaceous Interior seaway, and only thin shelf deposits accumulated there. The episode terminated with regional flooding and development of a Late Turonian condensed, maximum flooding horizon across the northern Gulf shelf, recording waning sediment supply and renewed subsidence (Figure 10). Condensation and/or erosion is also suggested by contact relationships with the overlying Coniacian strata from south Texas to north Louisiana (Lundquist, 2000). The Coniacian through Santonian was a time of global eustatic sea-level highstand. Depositional style changed dramatically in the northern GOM. The Austin depositional episode (Figure 10) is defined and named for the blanket of chalk that covered the northern Gulf (Lundquist, 2000). The northern Gulf was dominated by extensive deep carbonate shelves (Figure 18) that extended to and beyond present outcrop. Austin deposits are characterized by the chalks created from deposition of organic-rich globigerinid and coccolith oozes on a deep, clastic-starved shelf. Pelecypod and echinoderm-rich grainstones, mudstones, marls, and calcareous shales are also widespread. The northwest Gulf remained an open platform connecting to the Cretaceous Interior seaway (Figure 18). Currents flowing across the connection between the two large oceanic basins may have played a role in creation of the distinctive intraformational scours and hard grounds that typify the Austin chalk in northeast Texas (Hovorka and Nance, 1994). Interchange of the Boreal water mass of the Interior Seaway with Tethyan water mass of the Gulf is recorded by presence of mixed faunas in central Texas (Lundquist, 2000). A minor pulse of clastic sediment supply rebuilt local coastal deposits (Eutaw Formation) across the innermost northwestern and central shelf, but these were a faint ghost of the earlier Woodbine and Tuscaloosa fluvial/deltaic systems. Most of the clastics contain abundant glauconite and carbonate grains, reflecting extensive reworking in shore-zone and shallow-shelf environments associated with ongoing transgressive flooding of older coastal plain deposits during the Austin episode. Mud cracks and intertidal features indicate local carbonate shore-zone deposition above the San Marcos arch (Figure 18). The Cretaceous shelf margin-foundered and was blanketed by a ramp-like wedge of fine carbonate sediments. In contrast to the hundreds of meters of Austin strata found on the northern shelf, the deep, central Gulf was largely sediment-starved during this interval of regional highstand. Across the north-central shelf, from the East Texas basin to the Monroe uplift, tuffs and bentonites record extrusive volcanism from several vents located in southern Arkansas and on the Monroe uplift (Byerly, 1991).

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This volcanic activity, which continued through the Late Cretaceous, may be a residual effect of mantle plume activity (Cox and Van Arsdale, 2002). The Austin depositional episode, although characterized by accumulation of open-shelf carbonates across the northwestern Gulf, nonetheless records a shoaling cycle bounded by periods of relatively deep water (Lundquist, 2000). The boundary between the Austin and overlying Campanian Taylor episode is regionally disconformable. Updip basal Taylor strata in Arkansas contain abundant glauconite, phosphorite, shark teeth, and shells, typical of marine condensed and shelf deflation horizons. A deepening event separates mud- and limestone-dominated Lower Taylor from sandy Upper Taylor episode strata. Upper Taylor episode deposition (Figure 10) was characterized by renewed sandy terrigenous sediment influx to the Gulf margin, this time to depocenters in the northwestern part (Weise, 1980; Tyler and Ambrose, 1986; Sohl et al., 1991). Initially, muddy, sediment-laden plumes from southern Rocky Mountain-sourced delta and coastal systems of the southern Interior Seaway spread into the northwestern Gulf. By Late Campanian, the southern seaway had filled, and the wave-dominated San Miguel (Figure 9) delta system spilled across the remnant foreland basin into the Rio Grande embayment. This overflow of Laramide-sourced clastics created a depocenter that dominates the otherwise thin Campanian sequence of the Gulf. The mixing of Tethyan and Boreal water masses ceased, as the western Gulf again became an enclosed ocean basin. Additional siliciclastic material was locally provided by numerous volcanic cones that rose across the Rio Grande embayment and San Marcos arch, in South Texas, and over the Jackson dome (Byerly, 1991). Extrusion, intrusion, and crustal heating elevated the south Texas shelf, creating bioclastic grain shoals that constitute the Anacacho Limestone (Luttrell, 1977). Regionally across the central and northeastern Gulf, relative sea level remained high, submerging the basin margin throughout most of the Taylor depositional episode. Deposition occurred dominantly in shallow- to deep shelf systems. Even the fringing terrigenous deposits, found along the present outcrop belt, largely record shallow shelf, shoreface, and transgressive marine settings. Thus the genetic sequence consists of a mosaic of marine sediments including calcareous claystone, fossiliferous mudstone, glauconitic and fossiliferous sand, marl, chalk, and impure limestone. The terminal, Maastrichtian stage, depositional episode of the Cretaceous Gulf of Mexico is recorded by the Navarro Group (Figure 10). It too created a succession of strata that record a phase of siliciclastic-dominated progradation and shoaling bounded above and below by intervals of erosion, marine transgression, shelf starvation, and prominent flooding surfaces (Mancini and Puckett, 1995, 2005). In the northeast Gulf, shallow shelf sands, chalks, and marls bracket a Lower to Middle Maastrichtian shore-zone sand containing one or more inner-shelf disconformities (Skotnicki and King, 1989; Mancini and Puckett, 1995). Abundant lags of phosphorite, bored phosphatized mud clasts and fossil casts, turtle, shark, fish, and mosasaur teeth and bone fragments, and durable shell debris indicate nearshore to inner-shelf current erosion formed the disconformities,

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Paleogeography and principal depositional systems of the Upper Cretaceous Navarro depositional episode.

likely in response to relative sea-level fall. Collapse of the Mississippi embayment as cooling subsidence followed upon Late Cretaceous migration of the Bermuda hotspot eastward, had, by this time, created a marine reentrant that extended northward along the modern Mississippi Valley (Figure 19). In the east-central Gulf margin, progradation of the Nacatoch delta and shore-zone systems (Figure 19) records a significant clastic pulse (McGowen and Lopez, 1983). The Olmos delta system, the largest Navarro episode delta, prograded across the Rio Grande embayment from Laramide uplands in northern Mexico (Tyler and Ambrose, 1986). Several unconformities within and at the base and top of the Navarro Group record continued influence of Laramide crustal stresses on local uplift and subsidence across the northern Gulf basin (Tyler and Ambrose, 1986; Sohl et al., 1991). In general, maximum deltaic and shore-zone progradation occurred by late Middle Maestrichtian. Subsequent transgression of the northern Gulf margin formed an extensive flooding surface; however, local tectonics and sediment supply pulses created an extended period of latest Cretaceous retrogradational and highstand deposition that is here included in the 6 Ma Navarro depositional episode. The Cretaceous–Tertiary boundary strata of the Gulf of Mexico constitute a condensed horizon, recording widespread sediment starvation throughout the area of preserved Cenozoic strata. They also record a cataclysm of global proportions, the Chicxulub meteorite impact event (Hildebrand et al., 1991). The Chicxulub crater is located beneath the Yucatan Platform, in the southern Gulf. The impact crater forms an oval feature that is 90 by 120 miles (140 by 190 km) in diameter. The consequent seismic shock triggered submarine slides and mass flows (Bralower et al., 1998). An impact tsunami created a distinct event bed widely noted around the northern Gulf margin (Schulte et al., 2006).

5.4. Cenozoic depositional episodes The Cenozoic depositional history of the northern Gulf basin has been synthesized by Galloway et al. (1991a, 2000). Galloway et al. (2000) differentiated 18 northern GOM depositional episodes. Here, I have grouped these into 13 episodes (Figure 7) by combining some minor episodes and emphasizing only the first-order changes in supply history and paleogeography. These episodes can be further grouped into four families that record major evolutionary phases in the adjacent North American drainage basins. (1) Paleocene–Middle Eocene Laramide compression-related episodes. (2) Late Eocene–Oligocene episodes initiated by crustal heating, uplift, and volcanism in the southwestern United States and Mexico. (3) Miocene episodes that record erosional rejuvenation of eastern North American uplands. (4) Early Pliocene–Quaternary episodes that record rejuvenation of western interior drainage basins due to uplift, climate deterioration, and high-amplitude, high-frequency glacioeustatic sea-level change.

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Deposits of each episode are separated by regional transgressive marine shale tongues that contain at or near their base, a maximum flooding surface. These Cenozoic depositional episodes create the archetypal genetic stratigraphic sequences (Galloway, 1989b).

5.5. Laramide depositional episodes Regional flooding of the Gulf margin to and beyond present outcrop that terminated the Cretaceous persisted for the first few million years ago of the Paleocene. Widespread shelf mudrocks and marls of the Midway and Porters Creek Formations blanketed the northern Gulf margin. However, beginning in the Late Paleocene and Early Eocene, depositional outbuilding of the coastal plain, spearheaded by large delta systems centered in the Houston, Mississippi, and Rio Grande embayments, heralded the onset of successive waves of Cenozoic clastic influx (Galloway et al., 2000). Four principal depositional episodes punctuate Paleocene through Early Eocene history throughout the northern Gulf of Mexico (Figures 11 and 12). They record surges of clastic supply, modulated the progressive advance of Laramide uplift that began in the Central and Southern Rocky Mountains of the United States and spread progressively south to the Sierra Madre Oriental of northern Mexico (Winker, 1982; Galloway, 2005b). Laramide compressional crustal stress extended eastward into the Gulf basin, as reflected by broad folding in the Rio Grande embayment, rejuvenation of the Sabine and Monroe uplifts, and accentuated subsidence of the western Gulf abyssal plain (Laubach and Jackson, 1990; Cao et al., 1993; Feng et al., 1994). The Late Paleocene and Early Eocene Wilcox episodes significantly prograded the northern Gulf shelf margin and continental slope from its Cretaceous position (Figure 6B). The Lower Wilcox depositional episode records the first major Cenozoic influx of sediment onto the northern Gulf continental margin. A broad fluvial-dominated delta system prograded across the Houston embayment and onto the relict Cretaceous slope (Figure 6B). A second, smaller fluvial-dominated delta built across the Mississippi salt basin. Both form primary Late Paleocene depocenters. An extensive wave-built shorezone system extended across the San Marcos arch into northern Mexico. Rapid sediment loading mobilized the deep-water muds and Louann salt, initiating numerous extensional growth faults along the paleo-shelf margin. These growth faults form the inboard elements of the Wilcox fault zone, which extends from northern Mexico to central Louisiana (Figure 4). Loading also initiated the first of successive Cenozoic phases of salt mobilization and expulsion from beneath the basin-margin depocenters toward the paleo-continental slope, where salt canopies were initiated. In the northwestern Gulf, contemporaneous Laramide compression uplifted and tilted the underlying Cretaceous shelf deposits, which formed the foundation beneath Early Paleocene strata. Tilting triggered the Lobo megaslide, centered above the Rio Grande embayment, which affected more than 5,000 km2 of the Gulf margin. Burial of the megaslide created the third of the Lower Wilcox depocenters. Ongoing seismicity associated with foreland deformation of the West Gulf margin triggered frequent smaller slumps and slides along the prograding clastic shelf margin from south Texas to central Louisiana. Several of these slumps nucleated submarine canyons that excavated up to several hundreds of meters of older Wilcox strata (Galloway et al., 1991b). The Lower Wilcox depositional episode terminated with backstepping of delta and shore-zone facies. The Middle Wilcox, which was differentiated as a minor episode bracketed by two widely correlated, thin, marine shale horizons by Galloway et al. (2000), is here grouped with the Lower Wilcox. The transgressions, recorded by the regional Yoakum Shale and Big Shale markers, that punctuated the Paleocene-to-Eocene transition are associated with large submarine canyons. The best known canyon, the Yoakum, is located above the San Marcos arch in the central Texas coastal plain, cut across the transgressive shelf more than 150 km landward from the shelf edge, and excavated as much as 1.5 km of underlying Lower Wilcox deltaic deposits (Galloway et al., 1991b). A canyon of this size was not seen again on the northern Gulf margin until the Pleistocene (Galloway, 2005a). Following the Middle Wilcox ‘‘breather’’ in clastic supply and consequent transgression, which is reflected by the fossiliferous, glauconitic Sabinetown Formation in Texas outcrops and the Yoakum Shale in the subsurface, rejuvenated and reorganized Early Eocene bedload-dominated fluvial systems spilled across the San Marcos arch (Figure 20). The fluvial systems deposited an amalgamated network of coarse, sandy channel fills across the middle and south Texas coastal plain, creating the Carrizo aquifer, one of the major aquifer systems of the Gulf basin (Hamlin, 1988). Basinward, these fluvial systems supported a family of wave-dominated deltas that prograded rapidly to and over the shelf margin (Edwards, 1981). Here, basinward rafting of the underlying Mesozoic section opened an arcuate, growth-fault-bounded ‘‘depotrough’’ that collected highly expanded successions of delta front and slope apron sediment (Fiduk et al., 2004). Initial progradation into the head of the Yoakum canyon likely led to sediment bypass down the canyon and speculative formation of a submarine fan system at the base of the central Texas continental slope (Figure 20).

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Tectonic realignment of continental drainage systems diverted supply from the east-flowing tributaries that drained into the central Gulf during the Lower Wilcox episode. There, limited sediment supply and extensive marine reworking created broad, sandy shore-zone and shelf systems that trend northward along the flank of the deltaic coastal plain, which extended as far as east-central Louisiana, and into the Mississippi embayment (Figure 20). The Sabine uplift provided a low-relief upland drained by minor fluvial tributaries. The eastern Gulf of Mexico was largely a shallow shelf that, along with the basin floor, remained sediment-starved. Carbonate sediment accumulated throughout the length of the Florida platform. Beginning in the Paleocene, and continuing into the Eocene, the broad, moderately deep Suwanee strait (Figure 20) connected the northeast corner of the Gulf with the Atlantic Ocean. Strong marine currents were funneled through this strait, which separated siliciclastic and carbonate shelf provinces. By Upper Wilcox deposition, the northeast Gulf shelf had evolved the compound dip profile that is still reflected in the West Florida terrace and Florida escarpment (Figure 1). A shallow, perched, prograding shelf break, located near the present Florida coast line, separated the shallow clastic and carbonate shelf systems from a broad submarine ramp, which in turn was perched atop the foundered, relict Cretaceous deep shelf and fore-reef slope. The regional Recklaw transgression terminated the Wilcox depositional episode at about 49 Ma. Meanwhile, erosion and burial of Laramide southern Rocky Mountain uplands, which provided the principal source of sediment to the Gulf, resulted in diminishing sediment supply (Galloway and Williams, 1991). The Middle Eocene Queen City and Sparta episodes deposited sediment primarily on the Wilcox depositional platform (Figures 11 and 12). The continental slope and abyssal plain remained sediment-starved in the northern and eastern Gulf. The Queen City episode paleogeography resembled that of the Upper Wilcox. Deposition of wavedominated deltas and thick barrier and strandplain systems was centered in the northwestern Gulf, and an embayed, marine shelf extended across Louisiana and Mississippi and northward into the Mississippi embayment. The very broad, funnel-shaped embayment amplified the normally low tidal range of the Gulf and created, in Queen City deposits, a unique assemblage of tide-dominated shore-zone and shelf facies in the East Texas basin (Ramos and Galloway, 1990). Following the Weches transgression, the Sparta depositional episode records a shift of continental fluvial drainage axes back into the central Gulf, filling the Mississippi embayment with deposits of a fluvial-dominated delta system. The overall low rate of sediment supply and extensive but shallow marine flooding of the northern Gulf margin created widespread fossiliferous marine shale and glauconite beds that extend to outcrop and record long periods of very slow sediment accumulation. The Weches Formation (ca. 45 Ma), which separates deposits of the Queen City and Sparta episodes, is a muddy, fossiliferous glauconite sand that can be traced from northern Mexico to Mississippi, and records as much as 1 Myr of time in its few meters of sediment. The Cook Mountain transgression, which terminated the Sparta episode, also records about 2 Ma of northern Gulf coast inundation.

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5.6. Middle Cenozoic volcanism and related depositional episodes The latest Middle Eocene saw a modest rejuvenation of sediment influx onto the northwestern and central Gulf margin. Deposits of this episode, named the Yegua Group in Texas and Cockfield Formation in Louisiana, are also distinguished by a new appearance of volcanic ash beds. Initially, Yegua and Cockfield fluvial-dominated deltas prograded across the shallow transgressive shelf that had submerged Sparta delta and shore-zone systems in the Houston embayment and Mississippi salt basin. As the largest and most actively prograding Yegua deltas of the Houston embayment approached the shelf margin, they first built across the perched Sparta and Queen City delta platform margins, which formed a mid-shelf platform break. Progradation was then onto the much deeper, mudblanketed, distally steepened ramp that had evolved during more than 10 Myr of subsidence and tilting of the continental margin created by Upper Wilcox offlap. The combination of rapid sediment influx and renewed loading of the old, muddy continental margin triggered a succession of submarine slumps and growth faults that coalesced along the Yegua delta front to form a compound intraformational mass wasting surface that soles out within underlying muddy Eocene strata (Edwards, 1991). Slide scars extended as much as 20 km inland from the margin, creating steep slopes and local depocenters that both initiated and collected further mass flows and turbidity currents (Figure 21). Following this retrogradational phase, shelf-margin deltas built across the slump complex, healing the embayed margin and prograding the shelf edge. Sediment remobilized from the unstable shelf-margin delta front and prodelta formed a heterolithic slope apron. Yegua strata are well known for their excellent development of incised channels or valleys that extend many kilometers from platform delta lobes across a muddy outer shelf and terminate in small, low-stand, shelf-margin delta lobes. A series of transgressions and forced regressions created 5–10 (depending on location and author) significant progradational pulses during the 2.5 Ma Yegua episode (Edwards, 1991; Meckel and Galloway, 1996). High-effort micropaleontological analysis (Fang, 2000) of mid- and down-dip Yegua strata confirm that maximum flooding surfaces found within the bounding transgressive marine shelf mudstones are disconformities. In contrast, despite clear forced regression and channel cutting across shelf mudstones, no measurable hiatus could be documented across these surfaces within the Yegua genetic sequence. Additional delta systems prograded into the Rio Grande embayment and North Louisiana and Mississippi salt basins (Figure 21). Much delta sediment was reworked along strike to build thick, progradational, barrier and strandplain systems, particularly in the NW Gulf. The Suwannee strait was in the last stages of filling, but continued to separate the carbonate-dominated shelf of the Florida platform from the siliciclastic shelf and shorezone systems of Mississippi and Alabama. Eocene deposition terminated with the minor, but economically significant Jackson depositional episode (Figure 12). Following the tentative Yegua probe to the continental margin, which terminated with the regional Moodys Branch transgression of the northern Gulf, deposition during the next 2 Ma remained firmly on the

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up- to mid-dip platform (Figure 13) and sandy deltaic and shore-zone systems were restricted to the NW Gulf margin (Galloway et al., 1991a). Delta systems prograded only into the Houston embayment. Extensive barrier island and strandplain systems extended across the central and south Texas coast. For the last time, the Mississippi embayment suffered marine inundation as all of the central and NE Gulf margin reverted to an extensive muddy shelf. Jackson strata in Texas contain common bentonite and vitric ash beds, presaging the impending Oligocene climax of crustal heating and continental uplift. The Oligocene was a time of massive sediment influx to the Gulf (Galloway and Williams, 1991). The epoch began with extensive crustal heating, uplift, and volcanism of source areas in northern Mexico and the southwestern United States. Uplift impinged directly on the western margin of the Gulf Basin. Cretaceous and Early Cenozoic foreland basin fill was elevated more than 3 km in the Late Eocene through Early Oligocene along the SW Gulf margin (Gray et al., 2001). The NW margin, now the western edge of the Burgos basin, was similarly elevated in the Middle Oligocene. Further west, explosive volcanism and caldera collapse combined with the uplift to create a long-lived outpouring of recycled sedimentary rocks, volcaniclastics, and reworked, devitrified ash that peaked by the Mid-Oligocene and continued into the Early Miocene. The response in the Gulf was the sediment-supply dominated Frio depositional episode, which lasted for more than 8 Ma (Figure 12). Stratigraphic and structural architecture of the basal strata of the Frio episode, the Vicksburg Group, is complex. Uplift and volcanism directly affected the northwestern Gulf margin and indirectly affected the northcentral margin by rejuvenation of several drainage basin hinterlands and pervasive deposition of easily reworked ash. The preceding transgression of the Jackson coastal deposits was brief and most clearly recognizable in the shallow subsurface to mid-dip central Texas coastal plain (Galloway et al., 1994). At outcrop along the northwest Gulf margin, the boundary is variously manifested by the abrupt superposition of alluvial plain deposits on coastal Jackson facies, prominent mature paleosoils, locally inset basal Vicksburg alluvial channel and valley fill successions, and low-angle discordance between Jackson and basal Oligocene deposits (Galloway, 1977; Galloway et al., 1979; Combes, 1993). This assemblage of features shows that relative base level rise and transgression of the mid-dip Jackson fluvial and shore-zone systems were contemporaneous with mild tilting and relative uplift along the updip margin of the basin. Indeed, the beginning of the Oligocene marks a change in the style of tectonic subsidence along the Gulf margin (see Galloway et al., 1991a, Figure 3). Paleocene-Eocene subsidence involved minimal basinward tilting; sequences thicken only gradually until they reach the paleo-shelf margin. In contrast, Oligocene and all younger sequences thin rapidly as they approach their outcrop, indicating that tilting subsidence along a basinward progressing hinge has characterized the Late Cenozoic. The combined influences of continental uplift and concomitant deposition of massive amounts of air-fall ash in the various fluvial drainage basins is reflected in the total load, sediment composition and texture, and the progressive growth of the four delta systems that were active during Frio deposition (Galloway, 1977; Galloway et al., 1982b) (Figure 22). The primary Oligocene depocenter lies in the Rio Grande embayment and consists of up to 5 km of deposits of the Norias wave-dominated delta system and its associated fluvial and delta-fed apron systems (Galloway et al., 1982b). Norias deposition began with rapid progradation of the Vicksburg phase deltas onto a thick foundation of muddy Eocene shelf and slope deposits. The shelf margin was further destabilized by seismicity and uplift and tilting of the western Gulf margin. The immediate consequence was development of the Vicksburg detachment (Figure 5), a shale-based detachment system that extends more than 500 km along strike from the Burgos basin in northern Mexico (Diegel et al., 1995). This detachment created the Vicksburg growthfault zone, which forms the updip boundary of the much broader Frio fault zone (Figure 4). The shallow detachment within Upper Eocene mudstone resulted in horizontal displacement of Vicksburg delta front and upper slope deposits of as much as 16 km horizontally (Diegel et al., 1995). Following stabilization of this detachment, further Frio progradation built the continental margin 90–145 km beyond its Eocene position. Progradational loading of the continental slope initiated basinward advancing lines of growth faults that form the updip part of the northwest Gulf Oligo-Miocene detachment province (Figures 4 and 5). Basinward, extension was compensated by compressional faulting and folding in the Port Isabel fold belt, which lay at the base of the Oligocene continental slope (Figure 4). To the south, the Norma delta rapidly prograded into the Burgos Basin. Here, however, most tilting and uplift followed Early Oligocene progradation. A prominent, angular unconformity separates the Norma Conglomerate and equivalent ‘‘Non-marine Frio’’ from underlying Vicksburg and ‘‘Marine Frio’’. As uplift of the Sierra Madre Oriental migrated eastward, older Cenozoic strata were elevated and recycled. The third principal delta system, the Houston delta, is centered beneath the southeast Texas coastal plain (Figure 22). Initial Vicksburg delta lobes are thin, and largely remained on the Eocene shelf platform; growth faulting effected only their distal fringes. However, relative base-level fall along the inboard basin margin is indicated by incision of valley systems that extend from outcrop into the shallow subsurface (Galloway, 1977; Combes, 1993). As Frio deltas prograded into and across the Houston salt basin, loading of subjacent Louann salt fostered a phase of active salt diapir growth and minibasin development (Diegel et al., 1995). A fourth delta

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system, which was fed by a large, suspended load-rich fluvial system flowing south along the Mississippi embayment axis, prograded across the Louisiana salt basin (Figure 22). Here, large-scale salt evacuation from beneath the delta and continental slope depocenter accommodated as much as 4 km of Oligocene strata (Figure 6B). However, arrival of sediment to the central Gulf was delayed. Early Eocene Vicksburg strata of Louisiana consist of shelf mud and marl. The paleo-Mississippi fluvial system, unlike its sister systems of the northwestern and western Gulf, was not directly effected by uplift of its tributary drainage basin. Rather, sediment influx was accelerated by the rapid recycling of largely altered volcanic ash, in the form of suspended mud, through this mid-continental drainage system. Thus, the resultant delta system was large, but muddominated and slow to develop. Between the delta systems, the Frio sequence contains comparably thick successions of strandplain and barrier/lagoon complexes (Figure 22). These wave-dominated shore-zone systems were nourished by longshore reworking of sediment from the deltaic headlands. The central Texas barrier/lagoon complex contains as much as 1.5 km of stacked, amalgamated barrier, beach ridge, and shoreface sand (Galloway et al., 1982b). Together, the thick, prograding delta, shore-zone, and slope apron systems initiated and perpetuated a succession of growth faults that extend from northern Mexico to eastern Louisiana (Figure 4). Between the SE Texas and Louisiana delta systems, particularly rapid Mid-Oligocene salt withdrawal from beneath the shore-zone, shelf, and upper slope systems triggered a brief phase of tilting, collapse, and submarine erosion that interrupted margin progradation. The resultant Hackberry embayment is one of the best-described examples of many such destructional slope systems within the Gulf Cenozoic section (Cossey and Jacobs, 1992; Galloway, 1998a). The eastern Gulf basin remained clastic sediment starved. By Early Oligocene, the Suwanee strait had filled in, merging the Florida platform with the northeast shelf. Carbonate deposition on the outer shelf expanded westward as far as Louisiana. Local Late Oligocene patch reefs, known as the Heterostegina Limestone, developed over active salt domes in the Houston salt basin, the westernmost expansion of carbonate systems during the Cenozoic. Decreasing rate of sediment supply and accumulation in the Late Oligocene (Galloway and Williams, 1991) terminated the Frio depositional episode. Long-term backstepping of delta and shore-zone systems culminated in regional transgressive flooding and deposition of the Anahuac shale across the breadth of the Gulf margin.

5.7. Miocene depositional episodes Miocene basin fill reflects three multi-million-year depositional episodes that record the progressive shift of the locus of deposition in the Gulf of Mexico from the northwestern to the eastern margins. This shift reflects the

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concurrent reduction of the volcanic uplands, which sourced fluvial systems draining across northern Mexico and Texas, and the rejuvenation of the Appalachian and Cumberland Plateau uplands, which supplied rivers emptying into the central and east-central Gulf. Also during the Miocene, effects of Basin-and-Range tectonism extended to the western margin of the Gulf with activation of the Balcones fault system, uplift to the Edwards plateau, and development of the Rio Grande rift, which disrupted drainage into the Gulf from the southwestern United States. Consequent long-term changes in the rate and location of sediment supply largely defined three episodes that are approximately coincident with the Early, Middle, and Late Miocene (Figure 12). Concurrently, global climate was evolving toward the ice house world of the Late Cenozoic. Increasing amplitude and frequency of glacioeustatic sea-level fluctuations impacted stratigraphic and facies architecture, especially within deposits of Miocene shore-zone systems (Galloway, 1998b, 2002). The Gulf Miocene stratigraphy is characterized by extensive continental margin progradation (Figure 6B). By the Middle Miocene depositional episode, the dominant extrabasinal fluvial systems were established in positions that closely approximate Quaternary counterparts (Galloway et al., 2000; Galloway, 2005b). The Lower Miocene succession consists of an 8 Ma depositional episode that closely resembled major Paleogene episodes in its development (Galloway et al., 1986). An extended phase of high rates of sediment supply and continental margin outbuilding followed upon the Anahuac transgression. Following a transgressive interruption at about 18 Ma (used by Galloway et al. (2000) to differentiate two Lower Miocene sequences) that is best developed in the northwest Gulf, a 2 Ma phase of retrogradation and transgression terminated the episode. The widespread Amphistegina shale and its contained maximum flooding surface, which is named for the diagnostic Amphistegina B faunal top, caps the Lower Miocene genetic sequence (Figure 11). Following the Anahuac transgression, the bedload-dominated Rio Grande and Norma fluvial axes continued to decrease in relative importance, although they remained a major depocenter. Wave reworking and long-shore transport dominated the delta system, shifting the depocenter northeast to the laterally adjacent central Texas barrier-strandplain system. In the central Gulf, the paleo-Mississippi continued to increase in relative importance. A new fluvial axis, coincident with the modern Trinity/Sabine rivers, but with a drainage basin and size more commensurate with those of the modern Red River, entered the Gulf near the Texas/Louisiana border. Together, these two fluvial-dominated deltas prograded the continental margin 65–80 km basinward. At the onset of deposition, the Red and Mississippi deltas and slope aprons experienced a second episode of Hackberry-like hyper-subsidence and continental margin collapse and mass wasting. Numerous slump scars, fault-expanded shelf-margin deltas, and submarine canyon fills reflect the interplay of margin collapse, submarine erosion, and rapid deposition. The collapse of this ‘‘Planulina embayment’’ and concomitant development of the Planulina fault zone (Figure 4) was a consequence of large-scale salt withdrawal from beneath coastal Louisiana in response to the eastward migration of depositional loading. Combined deflation of the shallow, underlying Oligocene canopy and extension along the Oligocene and Louann detachment zones (Figure 8, panel C) (Diegel et al., 1995; Peel et al., 1995) accommodated nearly 7 km of Lower Miocene sediment in the central Gulf depocenter. Thick, sandy turbidite successions began to spill down the continental slope in the east-central and NE Gulf. Despite the proximity of the paleo-Mississippi delta system, the northeast Gulf margin initially remained a carbonate province. Much of the Alabama and Mississippi shelf was sediment-starved, creating a prominent nonconformity. Later in the episode, the delta-fed, muddy shelf encroached eastward, terminating reef growth and restricting carbonate platform deposition to the Florida shelf. Depositional loading beneath the deltas and well-nourished interdeltaic shore zones, and offlap of thick, sandy slope apron systems created prominent structural features, including the Lunker and Planulina fault zones (Figure 4). Compression continued along both the Port Isabel fold belt and initiated the Perdido fold belt, which lies along the basinward margin of the Louann salt. Salt extruded from beneath the prograding margin spread southward, nucleating new canopy complexes beneath the lower paleo-slope and adjacent abyssal plain. Inland along present outcrop belts in Texas, low-angle unconformities locally separate basal Miocene (Oakville Fm.) strata from underlying Oligocene (Catahoula Fm.) strata, and basal Middle Miocene (Goliad Fm.) strata from underlying Lower Miocene (Fleming Fm.) strata (Figure 12) (Galloway et al., 1982a, 1986). These discordances record intermittent tilting subsidence generated by successive episodes of sediment supply and crustal loading. The Middle Miocene sequence records a relatively brief (ca. 3 Ma) episode that was terminated by regional, but short-lived Gulf margin transgression associated with the Textularia stapperi faunal top. The paleogeography of the episode clearly documents the effects of Early Neogene continental tectonics and source area rejuvenation (Figure 23) (Galloway et al., 2000; Combellas-Bigott and Galloway, 2006). A new fluvial system, named for the Tennessee River, which currently occupies the comparable drainage basin, made its appearance. The system drained uplands characterized by Paleozoic outcrops and, consequently, transported sandy, mineralogically mature sediment to the Gulf. Together the paleo-Mississippi and paleo-Tennessee created the dominant Mid-Miocene

Depositional Evolution of the Gulf of Mexico Sedimentary Basin

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depocenter and prograded the continental margin as much as 70 km. Initial margin progradation was interrupted, however, by a third pulse of salt evacuation and hyper-subsidence, located beneath the southeast Louisiana coastal plain (Combellas-Bigott and Galloway, 2006). This ‘‘Harang embayment’’ and bounding fault zone (Figure 4) is the culmination of an Oligocene (Hackberry) to Miocene west-to-east wave of salt evacuation from beneath the prograding margin. Beneath the central Texas shelf, a second newly consolidated Corsair fluvial-deltaic system, prograded onto the continental slope (Morton et al., 1988; Galloway et al., 2000). Here, salt withdrawal and prolonged growth of the Corsair fault zone created a depocenter that was filled by wave-dominated delta and delta-fed apron deposits (Figure 23). Between deltaic headlands, extensive wave-dominated shore-zone systems were fronted by narrow, muddy to sandy shelves and prograding, muddy, shelf-fed slope aprons. In the northeast, combined margin collapse, slope bypass, and alignment of a series of dip-elongate slope minibasins created a relatively focused submarine transport pathway that diverted a large quantity of sediment from the paleo-Tennessee delta front to the adjacent slope toe and abyssal plain (Figure 23) (Combellas-Bigott and Galloway, 2006). The McAVLU submarine fan system (named for the three U.S. Minerals Management Service (MMS) protraction areas beneath which it lies) was born. It would persist as a major depositional feature of the eastern Gulf basin floor until the end of the Miocene. This and subsequent Neogene fan systems are distinguished from slope aprons by (1) location of a depocenter at the base of the contemporaneous continental slope, on the abyssal plain, (2) aggradational, rather than offlap, stratigraphic architecture, and (3) development of a radial sediment dispersal pattern indicating focused down-slope transport as a point source rather than a line source. By these criteria, fan systems are unusual features of the Gulf deep water; slope aprons are much more common and volumetrically important. Combined depositional loading and extension along the Gulf shelf margin caused continued compression along the Port Isabel and Perdido fold belts, triggered further shallow salt canopy inflation beneath the paleocontinental slope of Louisiana, and initiated the Mississippi fan fold belt on the northeast Gulf abyssal plain (Figure 23). Like the Perdido fold belt, compression along the Mississippi fan system was localized along the basinward pinchout of deep Jurassic salt. The Upper Miocene depositional episode records a long period (6 Ma) of relative paleogeographic stability and high sediment supply (Morton et al., 1988; Galloway et al., 2000; Wu and Galloway, 2002). Sediment input was dominated by the paleo-Mississippi and paleo-Tennessee systems. A large, compound fluvial-dominated delta system prograded onto the slope. Continental margin offlap occurred dominantly in the central Gulf, where the shelf edge advanced 40–90 km. The McAVLU fan continued to expand and evolve until late in the episode. However, to the west, the Corsair delta and surrounding shore-zone systems decreased in importance as sediment repositories. Wave reworking created an extensive strandplain, interrupted by several small wave-dominated deltas, from northern Mexico to eastern Louisiana.

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Depositional loading of the basin margin in the east-central Gulf by up to 5 km of Upper Miocene sediment continued to drive wholesale basinward salt displacement beneath the paleo-continental slope and abyssal plain and compression along the Mississippi fan fold belt. Along the curvilinear, wave-dominated northwest Gulf margin, continental margin progradation onto muddy slope aprons built the shelf edge to or near its present position. Here, loading created a linear belt of growth faults known as the Wanda fault zone (Figure 4). Compensatory contraction was focused along the northeastern segment of the Miocene compression domain (Figure 7B). The Upper Miocene episode terminated with regional marine flooding associated with the last occurrence of benthic foraminifer Robulus E and/or Bigenerina A. The subsequent 2 Ma depositional episode, which is named for the contained Buliminella 1 fauna, bridges the latest Miocene to Early Pliocene (Figure 12). Although sediment supply rates remained high and clastic input continued to be focused through the paleoMississippi and -Tennessee rivers, accumulation shifted back onto the continental shelf and margin (Galloway et al., 2000). Thickest deposits occur within a combined fluvial-dominated delta system and upper slope delta-fed apron on the central Gulf margin. Upper slope minibasins captured the bulk of the sediment that spilled over the shelf edge. Continued remobilization of the subjacent salt canopy is recorded in the South Timbalier Ship Shoal fault family, which is part of a larger roho domain (Figures 4 and 5) (Schuster, 1995). The middle and deep slope was under-girded by the extensive Miocene salt canopy complex. The McAVLU fan system was completely abandoned.

5.8. Early Pliocene–Quaternary depositional episodes Sediment influx and depositional patterns record the combined effect of further intracratonic tectonism, pronounced global and North American climate change, and high-frequency and amplitude glacioeustasy. As in the earlier Cenozoic, accumulation was concentrated along the continental margin and slope where depositional loading and salt migration produced the mosaic of minibasins and salt-cored highs. These minibasins have progressively been filled by advancing delta-fed aprons (Prather et al., 1998). Rapid, high-amplitude glacioeustatic sea-level changes are manifested in the Gulf stratigraphic record by development of multiple sequences of one to several hundred thousand years ago duration with well-defined subaerial exposure and flooding surfaces (Lawless et al., 1997; Weimer et al., 1998). Depositional paleogeography (Figures 24 and 25) and supply rate suggest these can be grouped into two low-order genetic sequences of about 2 Ma duration (Figures 11 and 12). 100°

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Following the brief, post-Buliminella 1 transgression, the pattern of deposition changed in several ways (Galloway et al., 2000; Galloway, 2005b): 1. The paleo Red River fluvial axis was rejuvenated. This reflects a response of its drainage basin to epeirogenic uplift and eastward tilting of the western High Plains and Rocky Mountains. 2. Sediment supply through the paleo-Tennessee continued to decline. As a consequence depocenters shifted to the northwestern Gulf margin, and the northeastern Gulf continental slope again became relatively sedimentstarved. There, Pliocene strata are thin, but sandy. 3. Shelf margin progradation occurred primarily along the west-central Gulf margin. 4. The northeastern upper slope and shelf edge locally retreated by combined subsidence and mass wasting, particularly in the Mid-Pliocene. Further evidence of slope instability is reflected in the development of a megaslide scar along the west flank of the delta system and correlative debris apron on the west Gulf abyssal plain (Figure 24). 5. Along the relatively steep northeast margin, turbidite channel complexes extended to the slope toe, initiating a new submarine fan system. This fan has been informally called the WRLU fan for its location beneath Walker Ridge and Lund MMS areas (Figure 24). Deposition in this fan system continued for much of the remaining Pliocene. 6. Across most of the slope, depositional loading of the shallow salt canopy began a process of molding the minibasin province that is reflected in the slope structure and topography of today (Figures 1 and 6). Beneath the outer shelf, salt withdrawal caused active growth of the South Cameron fault family (Figure 4). Oxygen isotopic data indicate inflow of glacial meltwater into the Gulf by latest Pliocene (Joyce et al., 1993). Development of the North American ice sheet profoundly altered drainage systems flowing into the Gulf. The paleo-Mississippi drainage basin was integrated as north-flowing streams were dammed and diverted south. Recurrent climate changes and consequent meltwater pulses began the process of excavation of the Mississippi Valley (Saucier, 1994) (Figure 1). As the valley was cut and back filled by glacial outwash, the Red and Tennessee Rivers were intermittently and then permanently trapped. The single Mississippi ‘‘Father of Waters’’ that now drains the middle of the United States was established by Late Pleistocene. Development of a singularly large river draining into the central Gulf of Mexico created an extensive fluvialdominated delta and subjacent slope apron system (Figure 25). At the same time, the high-amplitude, highfrequency glacioeustatic sea-level changes of the Pleistocene punctuated Pleistocene stratigraphy. Rapid transgressions forced shorelines temporarily landward 150–250 km, creating broad shelves. Subsequent sea-level draw downs carved deep valleys across the shelves and, together with the high rates of sediment supply, forced delta lobes to the shelf edge. Instabilities associated with rapid shelf edge deposition, pulses of glacial outwash and

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frequent sea-level changes triggered a phase of mass wasting and submarine canyon erosion and filling unlike any previously seen in the basin (Figure 25). Canyon excavation was most active on the east flank of the delta system. The Quaternary Mississippi fan system, the third in the succession of Neogene abyssal fan systems, has been fed through these canyons. Smaller, relatively short-lived canyons have created smaller fans, such as the Bryant fan. Beneath the prograding slope apron, minibasins continued to subside and fill as many delta-fed turbidite channel/lobe complexes and debris flows spilled down slope from prograding shelf-margin deltas. Salt mobilization and loading beneath the outer shelf is recorded by growth of the south Cameron and South Eugene Island fault families (Figure 4). The modern Gulf margin reflects, in sediment distribution and morphology, the latest Pleistocene Wisconsin lowstand and subsequent Holocene transgression. Much of the modern shoreline is relatively stable, lying at or near shoreline positions of the previous interglacial highstands. However, the Louisiana coastal zone is a product of the extensive Holocene progradation and abandonment of a succession of Mississippi delta lobes. Ongoing subsidence and wetland loss largely reflects the natural instabilities of such a young deltaic coast line.

6. Patterns and Generalizations in Gulf Depositional History The northern Gulf of Mexico history is long and complex. However, some common themes emerge. In addition, the high rate of sediment supply and accumulation creates an unusually complete record of nearly 160 Ma of North American geologic history.

6.1. Sediment supply: Sources and drainage history Siliciclastic depocenters and reconstructed paleogeographies of the northern Gulf margin reflect and amplify the story of tectonism within the North American continent. Mesozoic drainage axes primarily focused into the Maverick basin, East Texas basin, Mississippi salt basin, and Apalachicola embayment (Figure 26). They record five principal phases of continental drainage basin integration or reorganization: 1. Initially, syn-drift Jurassic fluvial systems arose in the remnant uplands of the southern Appalachian Mountains, and entered the northeast Gulf along the adjacent crustal sags (Figure 14). 2. By the Late Jurassic Cotton Valley episode, a southeastward-flowing fluvial system, arising from tributaries draining uplands in Colorado and New Mexico, created a new clastic depocenter within the East Texas basin (Figure 26, ETW axis). 3. High rates of clastic influx from multiple peripheral source areas were further augmented by Early Cretaceous intracontinental and basin-margin uplift at the termination of sea-floor spreading (creating the break-up unconformity and superjacent Hosston clastic episode). 4. Cenomanian thermal uplift of the Mississippi embayment, augmented by Laramide elevation of the Sabine and Monroe uplifts, rejuvenated fluvial systems of the Woodbine/Tuscaloosa episode (Figure 26, ETE axis). Principal rivers flowed into the East Texas and Mississippi salt basins. 5. Southward migration of Laramide uplift filled the southern remnant of the Cretaceous foreland basin, and clastics began to spill over into the Maverick Basin by the Late Campanian. During the Cenozoic era, five principal and three secondary, long-lived, extrabasinal fluvial/deltaic axes provided the bulk of the sediment that infilled the northern Gulf basin (Figure 27). Four major phases of continental uplift and erosion are recorded in the shifting patterns and rates of supply (Galloway, 2005b): 1. Palaeocene through Middle Eocene pulses of Laramide uplift along the Central and Southern Rockies and Sierra Madre Oriental supported the Early Cenozoic depositional episodes. Drainages focused through the cz, HN, and MS axes (Figure 27). 2. Late Eocene through Early Oligocene crustal heating, volcanism, and consequent uplift and erosion of much of central Mexico and the southwestern United States nourished major Oligocene through Early Miocene depositional episodes. The no, RG, HN, and MS axes (Figure 27) dominated input. 3. Initiation of erosion during the Early to Middle Miocene of the Cumberland Plateau and Appalachians invigorated supply to the east-central Gulf basin. At the same time, the high-standing Rocky Mountain uplands experienced continued regional exhumation. Whether climate change or uplift triggered this Miocene phase of erosion remains controversial; current literature favours climatic causes. In either case, sediment supply was concentrated in the cr, MS, and TN axes (Figure 27).

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Pliocene uplift of the western High Plains further rejuvenated northwestern sources and created a broad eastward slope that converged with the west-sloping alluvial apron of the eastern interior. Converging streams were variously combined and directed southward, forming the distinct Red, Mississippi, and Tennessee fluvial axes that dominated Middle Pliocene through Holocene sedimentation. High rates of Pleistocene sediment accumulation reflect rapid Quaternary climate cycling, and glacial erosion and runoff directly into the principal sediment transport systems. Only in the Late Pleistocene was the Mississippi valley sufficiently incised that the Tennessee and Red Rivers became permanently trapped within it (Saucier, 1994).

6.2. Climate and oceanography Climate setting of the Gulf basin has remained relatively constant throughout its history. The Gulf has generally lain within warm, subtropical climate zones. The Jurassic aridity of south-central North America is reflected in the widespread occurrence of evaporite, eolian, and sabkha deposits in the Louann and Smackover episodes. Evaporite deposits occur in Lower Cretaceous strata from northern Mexico to the Florida platform, suggesting continued hot, dry conditions. Late Cretaceous continental flooding likely led to a more equitable climate across the northern Gulf; however, limited preservation of terrestrial strata may also bias the preserved record of continental climate indicators. By Early Cenozoic, the climate of the northern Gulf margin was uniformly wet and tropical. Lignite deposits occur widely within Paleocene and Eocene fluvial, deltaic, and shore-zone systems. A dramatic climate change occurred at the beginning of the Oligocene Frio episode. Lignite deposition ceased. Carbonate-bearing paleosoils across the Texas coastal plain indicate rapid development of a strong east-west climatic gradient from wet subtropical in Florida to arid in northern Mexico (Galloway, 1977). This strong zonal pattern persists today.

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Deep Gulf oceanography experienced several major evolutionary milestones. The earliest (Middle Jurassic) opening was to the Pacific Ocean across Mexico (Salvador, 1991a). By onset of the Smackover episode, the Gulf had opened to the Atlantic. Through Late Jurassic deposition, the basin evolved into a small, east-west-elongate ocean basin open to the Atlantic through broad straits between the Yucatan and Florida platforms (Marton and Buffler, 1999). During the Albian, transgression and continental flooding established a shallow connection across northwest Texas with the southern end of the Cretaceous Interior seaway. The connection persisted until the Campanian. Strong marine currents periodically flowed through this broad strait (Figure 18). The connection also allowed mixing of equatorial Tethyan faunas with cool-water faunas of the Boreal Cretaceous seaway (Lundquist, 2000). Following the re-emergence of the basin margin following the Late Cretaceous flooding, a new pathway to the Atlantic, the Suwannee strait (also known as the Gulf trough) extended from the northeast Gulf across Georgia to the Atlantic (Popenoe et al., 1987). Initially, this strait formed a relatively deep trough that limited southward diffusion of terrigenous clastic sediment southward onto the Florida platform (Figure 20). Large erosional scours along the southeast Atlantic coastal plain indicate strong currents flowed through the trough. As Eocene deposition progressed, the strait shoaled and filled. Bridging of the strait is recorded in the appearance of siliciclastic beds within the previously pure carbonate platform facies of south Florida (Missimer and Ginsburg, 1998; Guertin et al., 2000). In the Middle Miocene, the first appearance of strong, deep-marine currents flowing through the Florida and Yucatan straits is recorded by erosion on the Florida escarpment and outer platform (Mullins et al., 1988; Guertin et al., 2000) and the first appearance of contourite drift deposits in the western Gulf abyssal plain (Galloway et al., 2000). This current system persists today as the Loop current and related systems. Glacial outflow from the North American Laurentide ice sheet first entered the Gulf in the Late Pliocene. The Pleistocene depositional episode as defined here extends to about 2.2 Ma to incorporate this milestone in northern Gulf margin deposition.

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The broad east-west fetch and prevailing west-to-east winds over the Gulf, at least through the Cenozoic, is reflected by the dominance of strandplain, barrier island, and wave-dominated delta systems along the north-west Gulf margin (Figures 20–25) (Galloway et al., 2000). However, the volume of sediment actually preserved in shore-zone systems relative to delta systems shows a pronounced and inexorable decrease beginning in the Middle to Late Miocene, coincident with increasing frequency and amplitude of glacioeustasy (Galloway, 2002). Development of the strong east-west climate gradient in the Oligocene is coincident with a period, extending from latest Eocene through Middle Miocene, which shows greatest development of shore-zone systems. Tidal modification of northern Gulf of Mexico coastlines has remained minimal. Tidal facies formed in local environments, such as barrier inlets and bays, that enhanced the typically microtidal range. Only at a few times and locations did an ideal combination of shelf width and embayed coastal geography lead to development of regionally tide-dominated coasts (e.g., Queen City episode; Ramos and Galloway, 1990).

6.3. Continental margin evolution The continental margin of the northern Gulf of Mexico experienced five general phases in its development (Figure 28A) (Winker and Buffler, 1988). 1. Evolution from ramp to defined, prograding clastic shelf–slope break during the Smackover through Hosston depositional episodes. 2. Stabilization of a rimmed carbonate platform and accentuation of slope-to-basin relief during the Early Cretaceous Sligo through Washita episodes. 3. Brief progradation of perched shelf-margin deltas and slope apron during the Woodbine/Tuscaloosa episode. 4. Drowning and blanketing by deep shelf and ramp deposits of the Late Cretaceous episodes. 5. Cenozoic offlap. Continental margin outbuilding has primarily been accomplished during clastic-dominated episodes by progradation of large delta systems and, as a well-developed slope developed, of their subjacent, sandy slope aprons (Prather, 2000; Winker and Booth, 2000; Galloway, 2005a) (Figure 28B). Between deltaic headlands, along-strike reworked sediment of shore-zone and shelf systems also spilled over the shelf edge, creating subordinate, but also extensive shelf-fed aprons, which consist largely of muddy sediment. In the northeastern Gulf, Cretaceous shelf margins retreated up to several hundred kilometers landward in response to subsidence and long-term sea-level rise. Subsequent Cenozoic margins have advanced through a combination of deposition and global sea-level fall. Although dominantly depositional in origin, and reflecting the long-term domination of sediment supply over subsidence, the Cenozoic continental margins of the northern Gulf record numerous phases of shelf edge and slope retreat and erosion (Edwards, 2000; Galloway et al., 2000) (Figure 28C). Such margins can be considered as destructional (Galloway, 1998a), in comparison to the offlapping, constructional continental margins that dominated Gulf deposition. Gulf destructional margins form three general groups, depending on their morphology and origin. 1. Submarine canyons are dip-elongate, erosional troughs that are hundreds of meters deep. Canyons generally occur in middle- to upper slope strata, but may extend onto the shelf as much as several tens of kilometers. They were excavated by combined processes of submarine mass wasting and turbidity current flow. Large submarine canyons cluster along margins constructed by the Paleocene–Early Eocene Wilcox depositional episodes (Figure 28C, features 4–7) and in the Late Pliocene Pleistocene margins (Figure 28C, features 18–21). 2. Megaslides are laterally extensive features that include extensive slump or slide deposits and an embayment within the upper slope and shelf edge bounded by a prominent slump scar and glide plane faults. Examples include features 2, 3, 15, and 17 (Figure 28C). They record brief, catastrophic failures of the continental margin due to tectonic tilting, salt withdrawal, or sedimentary loading. The latest Cretaceous slide (feature 2) is the most exotic if it was, in fact, triggered by the Chicxulub meteorite impact event. 3. Retrogradational slopes are the largest, longest lived, and most complex destructional slope type. Examples occur along the Yegua (feature 10), Frio (feature 12), Lower Miocene (feature 13), Middle Miocene (feature 14), and Pliocene (feature 16) margins. Such slopes exhibit complex mixtures of slump scars, canyons and erosional channels, growth faults, and thick sections of mixed shelf-margin delta, turbidite and debris flow facies. Most are distinguished by a muddy sediment wedge containing deep-water faunas and by concave reentrants in the shelf margin. Most appear to be related to pulses of evacuation of subjacent thick primary salt or salt canopies, causing subsidence and tilting and destabilization of a segment of the slope and outer shelf. Tilting led to oversteepening and slope failure. Forming and filling of the embayments commonly required more than a million years.

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Figure 28 Mesozoic and Cenozoic shelf margins and related features of the Northern Gulf Basin. Mesozoic: JS, Smackover, LK, Hosston; KS, Sligo; KWF,Washita-Fredericksburg (Stuart City reef); UK, Upper Cretaceous (Tuscaloosa/Woodbine); LW, Lower Wilcox; UW, Upper Wilcox; QC, Queen City; Y,Yegua; JS, Jackson; F, Frio; LM, Lower Miocene; MM, Middle Miocene; UM, Upper Miocene; PB1, Buliminella 1; Plio, Pliocene; Pleist, Pleistocene. (A) Shelf margins at maximum o¥ap of principal depositional episodes. Modi¢ed from Galloway (2005a). (B) Relationship between shelf margin advance and location of major shelf-margin deltaic depocenters. (C) Location of principal submarine canyons, slides, slump scars, and compound retrogradational slope complexes. 1, Cretaceous canyon; 2,Top Cretaceous megaslide; 3, Lobo megaslide; 4, Lavaca/Smothers canyons; 5,Yoakum canyon; 6, Hardin canyon; 7, St. Landry canyon; 8, Upper Wilcox slumps; 9, Queen City slumps; 10, Lower Yegua retrogradational slope; 11, Lower to Middle Frio canyon; 12, Hackberry embayment/retrogradational slope; 13, Planulina embayment/retrogradational slope; 14, Harang embayment/retrogradational slope; 15, Middle Miocene megaslide; 16, Globoquadrina altispira (Middle Pliocene) retrogradational slope; 17, Globoquadrina altispira (Middle Pliocene) megaslide; 18, Late Pliocene slump and canyon cluster; 19, Early Pleistocene canyon cluster, 20, Late Pleistocene canyon cluster; 21, Bryant canyon; 22, DeSoto canyon; 23, Alabama scour trough.

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Depositional Evolution of the Gulf of Mexico Sedimentary Basin

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Figure 28 (Continued ).

Gulf submarine canyons, slides, and retrogradational embayments in the Gulf, as in many basins, have popularly been related to eustatic sea-level fall. Such an explanation does not explain their localization along otherwise normally prograding margins, their diversity, nor their common paleogeographic location on the margins of major deltaic depocenters. On the other hand, correlations with structural elements and events are clear for many of the Gulf features.

7. Energy Resources The Gulf of Mexico basin is a world class repository of hydrocarbons, and has produced significant lignite and sedimentary uranium (Nehring, 1991; Riggs et al., 1991). It has been actively explored for nearly 100 years, creating a three-dimensional well and reflection seismic data base of unique abundance, extent, and diversity. Hydrocarbon exploration and development moved off-shore in mid-century, and has progressed to the slope, and now onto the abyssal plain (Crawford et al., 2003). As stated by Nehring (1991), ‘‘No other basin worldwide has even come close to producing so many discoveries for such a long period’’. Because of this history, the northern Gulf has served, for more than 50 years, as a natural laboratory for understanding the sedimentary processes, facies, stratigraphy, and gravity tectonics of prograding continental margins. Unlike many basins, where a limited volume of the stratigraphic section produces the bulk of reserves, major hydrocarbon plays are found throughout the basin fill in reservoirs nearly every depositional episode. In descending volumetric rank, hydrocarbons occur in Miocene, Oligocene, Paleocene–Eocene, Plio-Pleistocene, Upper Cretaceous, Lower Cretaceous, and Jurassic strata. Total produced and known reserves of the U. S. part of the Gulf of Mexico basin exceed 150 BBOE (billion barrels of oil equivalent) (Nehring, 1991; Crawford et al., 2003). Of this total oil and condensate aggregate about 70 Bbl; natural gas volumes are about 500 Tcf. Ongoing exploration had discovered very large reserves of oil within Paleogene and Miocene reservoirs beneath the continental slope and in tight sand and shale gas reservoir systems within Jurassic and Early Cretaceous units (Cotton Valley, Bossier, Hosston) of the north-central Gulf margin. These will substantially increase ultimate recoverable recovery in the basin, and continue to make the Gulf basin one of the most active exploration theatres in the world. Thermally mature petroleum source rock strata occur in Jurassic, Cretaceous, and Early Cenozoic basinal marls and shales, principally deposited during Eocene, Turonian, and Tithonian–Oxfordian depositional episodes (Hood et al., 2002). Generation phases have extended over several tens of millions of years, depending on source level, burial history, and ambient heat flow. A hallmark of Gulf petroleum systems is large-scale vertical migration from Mesozoic source rocks into Cenozoic reservoirs.

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The long and varied history of deposition and immense volume of porous rock, combined with the equally long and complex history of gravity deformation and salt migration have created a diversity of reservoir and trap combinations rarely matched in global petroleum mega-provinces. Stratigraphic, structural, and combined trap types abound. The array of structures created by salt deformation and migration have played a particularly important role in trap development. In addition to its hydrocarbon wealth, the Gulf of Mexico basin contains large reserves of bituminous coal and lignite (Riggs et al., 1991). Lignite is extracted from Lower Wilcox, Yegua, and Jackson strata in Texas and Louisiana. Bituminous coal occurs in Upper Cretaceous strata of northern Mexico and along the Rio Grande River. Sedimentary uranium deposits occur along the South Texas coastal plain in strata of the Upper Wilcox (Carrizo Sandstone), Jackson (Whitsett Sandstone), Frio/Vicksburg (Catahoula Formation), Lower Miocene (Oakville Sandstone), and Middle Miocene (Goliad Sand) depositional episodes (Riggs et al., 1991). Uranium ore occurs along the irregular boundaries between reduced and oxidized parts of aquifer sand bodies known as roll fronts because of their ‘‘C’’-shaped cross-section. Uranium was leached from reworked air-fall ash associated with the Middle Cenozoic volcanogenic phase (principally within the Oligocene episode), transported by groundwater into the shallow fluvial and coastal sand aquifers, and trapped by reduction by detrital organic matter and/or epigenetic sulfide minerals. Ore has been mined both by open pit and in-situ solution methods.

ACKNOWLEDGMENTS Jeffrey Horowitz drafted the figures. Co-workers Patricia Ganey-Curry, Timothy Whiteaker, and Lisa Bingham aided in preparation of selected figures. Comments of reviewers Jim Dixon and Andrew Miall significantly improved the clarity and content. Angela McDonnell reviewed the final draft with care and efficacy.

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W., 1984, Cotton valley (Upper Jurassic) and Hosston (Lower Cretaceous) depositional systems and their influence on salt tectonics in the East Texas basin, The University of Texas at Austin, Bureau of Economic Geology Geological Circular 84–5. McGowen, M. K., and Lopez, C. M., 1983, Depositional systems in the nacatoch formation (Upper Cretaceous), northeast Texas and southwest Arkansas, The University of Texas at Austin, Bureau of Economic Geology Report of Investigations 137. Meckel, L. D., III, and Galloway, W. E., 1996, Formation of high-frequency sequences and their bounding surfaces: case study of the Eocene Yegua formation, Texas Gulf coast, USA. Sedimentary Geology, v. 102, pp. 155–186. Missimer, T. M., and Ginsburg, R. N., 1998, Homogenized carbonates and siliciclastics in the tertiary of southwest Florida. Gulf Coast Association of Geological Societies Transactions, v. 48, pp. 263–274. Moore, C. H., 1984, The Upper Smackover of the Gulf rim: depositional systems, diagenesis, porosity evolution, and hydrocarbon production, in Gulf Coast Section SEPM 3rd Annual Research Conference, pp. 283–308. Morton, R. A., and Ayers W. B. Jr., 1992, Plio-Pleistocene genetic sequences of the southwestern louisiana continental shelf and slope: geologic framework, sedimentary facies, and hydrocarbon distribution, The University of Texas at Austin, Bureau of Economic Geology Report of Investigations 210. Morton, R. A., Jirik, L. A., and Galloway, W. E., 1988, Middle-upper Miocene depositional sequences of the Texas coastal plain and continental shelf: geologic framework, sedimentary facies, and hydrocarbon plays, The University of Texas at Austin, Bureau of Economic Geology Report of Investigations 174. Mullins, H. T., Garduski, A. F., Hinchey, E. J., and Hine, A. C., 1988, The modern carbonate ramp slope of central west Florida. Journal of Sedimentary Petrology, v. 58, pp. 273–290. Murray, G. E., 1961, Geology of the Atlantic and Gulf coastal province of North America, New York, Harper and Brothers. Nehring, R. N., 1991, Oil and gas resources, in Salvador, A. ed., The Gulf of Mexico Basin, The Geology of North America, Geological Society of America, Boulder, CO, v. J, pp. 445–494. Nelson, T. H., 1991, Salt tectonics and listric-normal faulting, in Salvador, A. ed., The Gulf of Mexico Basin, The Geology of North America, Geological Society of America, Boulder, CO, v. J, pp. 73–89. Oliver, W. B., 1971, Depositional systems in the woodbine formation (Upper Cretaceous), northeast Texas, The University of Texas at Austin, Bureau of Economic Geology Report of Investigations 73. Peel, F. J., Travis, C. J., and Hossack, J. R., 1995, Genetic structural provinces and salt tectonics of the Cenozoic offshore U.S. Gulf of Mexico, in Jackson, M. P. A., Roberts, D. G., and Snelson, S. eds., Salt tectonics: a global perspective, American Association of Petroleum Geologists, Tulsa, OK, pp. 153–176. Pindell, J. L., and Kennan, L., 2001, Kinematic evolution of the Gulf of Mexico and Caribbean, in Gulf Coast Section SEPM 21st Annual Research Conference, pp. 193–220. Popenoe, P., Henry, V. J., and Idris, F. M., 1987, Gulf trough—the Atlantic connection. Geology, v. 15, pp. 327–332. Prather, B. E., 1992, Origin of dolostone reservoir rocks, smackover formation (Oxfordian), northeastern Gulf Coast, U.S.A. AAPG Bulletin, v. 76, pp. 133–163.

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Prather, B. E., 2000, Calibration and visualization of depositional process models for above-grade slopes: a case study from the Gulf of Mexico. Marine and Petroleum Geology, v. 17, pp. 619–638. Prather, B. E., Booth, J. R., Steffens, G. S., and Craig, P. A., 1998, Classification, lithologic calibration, and statigraphic succession of the seismic facies of intraslope basins, deep-water Gulf of Mexico. AAPG Bulletin, v. 82, pp. 701–728. Ramos, A., and Galloway, W. E., 1990, Facies and sand-body geometry of the queen city (Eocene) tide-dominated delta-margin embayment, northwest Gulf of Mexico basin. Sedimentology, v. 37, pp. 1079–1098. Riggs, S. R., Ellison, S. P., Jr., Fisher, W. L., Galloway, W. E., Jackson, M. L., and Morton, R. A., 1991, Mineral resources and geopressured energy, in Salvador, A. ed., The Gulf of Mexico Basin, The Geology of North America, Geological Society of America, Boulder, CO, v. J, pp. 495–528. Roesink, J. G., Weimer, P., and Bouroullec, R., 2004, Sequence stratigraphy of Miocene to Pleistocene sediments of east-central Mississippi canyon, northern Gulf of Mexico. Gulf Coast Association of Geological Societies Transactions, v. 54, pp. 587–601. Rowan, M. G., 1995, Structural styles and evolution of allochthonous salt, central Louisiana outer shelf and upper slope, in Jackson, M. P. A., Roberts, D. G., and Snelson, S. eds., Salt tectonics: a global perspective, American Association of Petroleum Geologists, Tulsa, OK, pp. 199–228. Rowan, M. G., and Weimer, P., 1998, Salt-sediment interaction, northern green canyon and Ewing bank (offshore Louisiana), northern Gulf of Mexico. AAPG Bulletin, v. 82(5B), pp. 1055–1082. Rowan, M. G., Inman, K. E., and Fiduk, J. C., 2005, Oligo-Miocene extension at the Louann level in the northern Gulf of Mexico: kinematic models and examples. Gulf Coast Association of Geological Societies Transactions, v. 55, pp. 725–732. Salvador, A., 1987, Late Triassic-Jurassic paleogeography and origin of Gulf of Mexico basin. AAPG Bulletin, v. 71, pp. 419–451. Salvador, A., 1991a, Triassic-Jurassic, in Salvador, A. ed., The Gulf of Mexico Basin, The Geology of North America, Geological Society of America, Boulder, CO, v. J, pp. 131–180. Salvador, A., 1991b, Origin and development of the Gulf of Mexico basin, in Salvador, A. ed., The Gulf of Mexico Basin, The Geology of North America, Geological Society of America, Boulder, CO, v. J, pp. 389–444. Saucier, R. T., 1994, Geomorphology and quaternary geologic history of the Lower Mississippi valley, Vicksburg, MS, Mississippi River Commission. Sawyer, D. S., Buffler, R. T., and Pilger, R. H., Jr., 1991, The crust under the Gulf of Mexico basin, in Salvador, A. ed., The Gulf of Mexico Basin, The Geology of North America, Geological Society of America, Boulder, CO, v. J, pp. 53–72. Schlager, W., and Camber, O., 1986, Submarine slopes: angles, drowning unconformities, and self-erosion of limestone escarpments. Geology, v. 14, pp. 762–765. Schulte, P., Speijer, R., Mai, H., and Kontny, A., 2006, The Cretaceous—Paleogene (K—P) boundary at Brazos, Texas: sequence stratigraphy, depositional events and the Chicxulub impact. Sedimentary Geology, v. 184, pp. 77–109. Schuster, D. C., 1995, Deformation of allochthonous salt and evolution of related salt-structural systems, eastern Louisiana Gulf coast, in Jackson, M. P. A., Roberts, D. G., and Snelson, S. eds., Salt tectonics: a global perspective, American Association of Petroleum Geologists, Tulsa, OK, pp. 177–198. Scott, R. W., 1990, Models and stratigraphy of mid-Cretaceous reef communities, Gulf of Mexico, Tulsa, OK, SEPM (Society for Sedimentary Geology) Concepts in Sedimentology and Paleontology. v. 2. Scott, R. W., 1993, Cretaceous carbonate platform, U.S. Gulf coast, in Simo, J. A. T., Scott, R. 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Weimer, P., Varnai, P., Budhijanto, F. M., Acosta, Z. M., Martinez, R. E., Navarro, A. F., Rowan, M. G., McBride, B. C., Villamil, T., Arango, C., Crews, J. R., and Pulham, A. J., 1998, Sequence stratigraphy of Pliocene and Pleistocene turbidite systems, northern Green Canyon and Ewing Bank (offshore Louisiana), northern Gulf of Mexico. AAPG Bulletin, v. 82 Part B, pp. 918–960. Weise, D. R., 1980, Wave-dominated delta systems of the Upper Cretaceous San Miguel formation, maverick basin, South Texas, University of Texas at Austin, Bureau of Economic Geology Report of Investigations No. 107, Austin, TX. Winker, C. D., 1982, Cenozoic shelf margins, northwestern Gulf of Mexico. Gulf Coast Association of Geological Societies Transactions, v. 32, pp. 427–448. Winker, C. D., 1984, Clastic shelf margins of the Post-Comanchean Gulf of Mexico: implications for deep-water sedimentation, in Gulf Coast Section SEPM 5th Annual Research Conference, pp. 109–120. Winker, C. D., and Booth, J. R., 2000, Sedimentary dynamics of the salt-dominated continental slope, Gulf of Mexico: integration of observations from the seafloor, near-surface, and deep subsurface, in Gulf Coast Section SEPM 20th Annual Research Conference, pp. 1059–1086. Winker, C. D., and Buffler, R. T., 1988, Paleogeographic evolution of early deep-water Gulf of Mexico and margins, Jurassic to Middle Cretaceous (Comanchean). AAPG Bulletin, v. 72, pp. 318–346. Worrall, D. M., and Snelson, S., 1989, Evolution of the northern Gulf of Mexico, with emphasis on Cenozoic growth faulting and the role of salt, in Bally, A. W. and Palmer, A. R. eds., The geology of North America—an overview. The Geology of North America, Geological Society of America, Boulder, CO, v. J, pp. 97–137. Wu, X., and Galloway, W. E., 2002, Upper Miocene depositional history of the central Gulf of Mexico basin. Gulf Coast Association of Geological Societies Transactions, v. 50, pp. 1019–1030. Wu, S., Vail, P. R., and Cramez, C., 1990, Allochthonous salt, structure and stratigraphy of the north-eastern Gulf of Mexico. Part I: stratigraphy. Marine and Petroleum Geology, v. 7, pp. 318–333. Yurewicz, D. A., Marler, T. B., Meyerholtz, K. A., and Siroky, F. X., 1993, Early Cretaceous carbonate platform, north rim of the Gulf of Mexico, Mississippi and Louisiana, in Simo, J. A. T., Scott, R. W., and Masse, J.-P. eds., Cretaceous carbonate Platforms, AAPG Memoir 56, Tulsa, OK, pp. 81–96.

CHAPTER 16

Geology of the Late Cretaceous to Cenozoic Beaufort-Mackenzie Basin, Canada Jim Dixon, J.R. Dietrich, L.S. Lane and D.H. McNeil

Contents 1. Introduction 1.1. Geographic setting 1.2. Previous work 2. Regional Geologic Setting 2.1. Tectonic elements 2.2. Plate tectonic setting 3. Stratigraphy and Sedimentology 3.1. Stratigraphy 3.2. Sedimentology 3.3. Stratigraphic history 4. Structure and Tectonics 4.1. Basin-margin faults 4.2. Beaufort Foldbelt 4.3. Listric faults 4.4. Timing of deformation 5. Economic Geology 6. Summary Acknowledgments References

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Abstract The Canadian Beaufort-Mackenzie Basin lies on the continental margin of the Arctic Ocean and contains 14–16 km of Late Cretaceous to Recent sediment, the bulk of which is Tertiary in age. It has rifted basin margins although the southeastern margin has been overprinted by compressional structures that are contemporaneous with deposition and are part of the Cordilleran foldbelt that extends along west North America. The combination of post-rift subsidence and tectonic loading explains the thick accumulation of Late Cretaceous to Recent sediments. These sediments have been divided into several large transgressive-regressive sequences, each one dominated by the formation of delta complexes and their lateral equivalents. Sediment was supplied mostly from the rising orogenic belt on the southwestern margin of the basin. Several phases of tectonism have been recognized, each terminated by a major basin-scale unconformity. The first phase extended from the Late Cretaceous to about the Middle Eocene, the second from Middle Eocene to the Late Miocene, and the third from the Late Miocene to the present. The first two phases are dominated by compressional tectonics, although listric faults and some translational faults also formed. The third phase is characterized by the lack of any significant deformation. The basin has significant reserves of oil and gas and its potential of additional resources is high. Exploration in the basin started in the 1960s but has undergone several periods of exploration slow-down. At present there is renewed activity and undoubtedly this basin will be a major source of hydrocarbons in the future.

Sedimentary Basins of the World, Volume 5 ISSN 1874-5997, DOI 10.1016/S1874-5997(08)00016-6

r 2008 Elsevier B.V. All rights reserved.

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1. Introduction The Beaufort-Mackenzie Basin underlies the Canadian part of the Beaufort Sea on the southeast margin of Canada Basin (part of the Arctic Ocean) and extends onshore for a short distance into northern Yukon and adjacent Northwest Territories (Figure 1). It is a Maastrichtian to Recent sedimentary basin that formed on a passive continental margin that has been overprinted by Tertiary compressional tectonics. We will describe the geographic, stratigraphic, and structural framework of the basin, its tectonic and depositional history, and briefly outline the economic importance of the basin. Much of our understanding of this basin is derived from the interpretation of reflection seismic data, and exploration and development wells (Figure 2). These types of data are available to the public through Canada’s National Energy Board. Paper copies of seismic data are made public after a five-year confidentiality period, and wells after two years. A limited amount of information about the Upper Cretaceous and Tertiary succession has been derived from onshore outcrops.

Figure 1

Location map and physiographic elements of the Beaufort-Mackenzie area.

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Figure 2 text.

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Location of wells: hydrocarbon-bearing wells/¢elds are identi¢ed, as are wells on cross-sections or named in the

1.1. Geographic setting The landward margin of the basin has three distinct physiographic elements: on the west are the mountains and plateaux of the Cordillera, with a narrow coastal plain, the centrally located Mackenzie Delta, and the lowlands of Tuktoyaktuk Peninsula and Anderson Plain on the east (Figure 1). Most of the basin underlies the shelf and slope of the Beaufort Sea. The shelf is up to 120 km wide offshore from the delta, but is as narrow as 10 km in the west Beaufort area. A northwest-oriented bathymetric trough, the Mackenzie Trough (Figure 1), located to the west of the delta, divides the shelf into two areas. The trough attains depths of up to 400 m before merging into the continental slope and at its southern end the 100 m isobath reaches within 10 km of the coastline. The depth of the shelf-slope break varies, from 60 m north of the delta, increasing northeastward to about 500 m off Amundsen Channel. The continental slope extends to water depths of about 1,100–1,800 m where it merges into the continental rise which extends for several hundred kilometers into Canada Basin.

1.2. Previous work Detailed studies of the Beaufort-Mackenzie Basin began with its onshore part, such as the mapping of Norris (1981a, 1981b, 1981c, 1981d) and the early stratigraphic studies of Mountjoy (1967), Chamney (1969, 1973), Young (1975), and Young et al. (1976). With increased interest in its hydrocarbon potential during the 1960s and into the 1980s, reflection seismic and well data became available, leading to a broader understanding of the basin (Lerand, 1973; Yorath, 1973; Hawkings and Hatlelid, 1975). As exploration expanded our knowledge, local and basin-scale studies began to be published (e.g., Dixon et al., 1984; Willumsen and Cote, 1982; Young and McNeil, 1984). During the late 1980s and into the 1990s, little exploration was conducted, which allowed time to synthesize the data, resulting in a number of comprehensive studies of the basin (Dixon et al., 1992, 1994; Dixon, 1995; Lane and Dietrich, 1995; McNeil, 1989; McNeil et al., 1990). These are by no means an exhaustive list of references but they constitute the major contributions over the past few decades and the reader can find more in the cited volumes, especially in Dixon (1995).

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Pre-mid Cretaceous strata have been examined in varying detail and the reader is referred to Dixon (1992a, 1995), Poulton et al. (1982), Wielans (1992), and the references therein.

2. Regional Geologic Setting 2.1. Tectonic elements The Beaufort-Mackenzie Basin formed on a post-rift, continental margin. Rifting began in the Jurassic and continued to the end of the Albian (Cretaceous) (Embry and Dixon, 1990). Although thermal subsidence began in the Cenomanian, thick accumulations of sediment did not begin to be deposited on the continental margin until the Late Maastrichtian. The bulk of Cenomanian to Middle Maastrichtian sediment was deposited on the continental crust to the south, in the Porcupine-Mackenzie Basin (Dixon et al., 1992). The southeast margin consists of a series of down-to-basin, crustal-scale listric faults. The innermost of these major faults defines the seaward margin of the Eskimo Lakes Arch (Figure 3). On the southwest margin, Late Cretaceous and Tertiary compressional overprinting masks the original rifted margin. In the northern Yukon a number of tectonic uplifts and depressions flank the southeast basin margin (Figure 3). Most of these features are Late Cretaceous to Tertiary in age although many are reactivated older tectonic elements. Within the basin there are four broad structural domains, distinguished by the style of deformation (Figures 3 and 4). Along the southwest margin, the basin is characterized by arcuate trending folds, commonly associated with north to northeast directed thrust faults (Figures 3 and 4a). Within this domain of arcuate folds are three tectonic elements of note, the Demarcation Sub-basin and associated Herschel High (Dietrich et al., 1989a; Dietrich and Lane, 1992), and the Blow River High. The central part of the basin, underlying Mackenzie Delta and the adjacent offshore, is characterized by north- to northwest-trending folds and generally east- to northeasttrending listric faults. Two prominent zones of large listric faults occur in this domain, the Taglu Fault Zone (TFZ) and the Tarsiut-Amauligak Fault Zone (TAFZ) (Figures 3 and 4b). In the far offshore of the western and central parts of the basin, east-trending folds with long axes are characteristic of this domain. The three domains

Figure 3

Simpli¢ed geological map and fold and fault trends in the Beaufort-Mackenzie area.

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Figure 4 Regional cross-sections, based on re£ection seismic, through the Beaufort-Mackenzie Basin: (a) western part of the basin, (b) central part of the basin (after Lane and Dietrich, 1995).

of folded and faulted strata are part of the Beaufort Foldbelt (Lane and Dietrich, 1995), which extends into adjacent Alaska. The final domain lies in the eastern part of the basin where there is virtually no folding and only minor faulting (Figure 3), and most of the faults are reactivated Mesozoic ones.

2.2. Plate tectonic setting The Arctic Ocean is one of the few remaining oceanic areas without an extensive database of aeromagnetic measurements and with limited ocean-bottom sampling. This is due in part to the extensive year-round sea-ice cover and the remoteness of the region, making access difficult and logistics expensive. Reconnaissance aeromagnetic data have revealed only a few areas of magnetic striping. In most of Canada Basin, magnetic stripes are masked by a thick sedimentary cover, or are overprinted by younger volcanics. Consequently the sense of relative motion of the adjacent continental plates is not well constrained. However, over the years a number of plate tectonic interpretations have been proposed (see Lawver and Scotese, 1990, for a review), most of which have not withstood close scrutiny. The most accepted interpretation is that of counter-clockwise rotation of northern Alaska away from the Arctic Islands, with a rotation pole within the northern Yukon (e.g., Rowley and Lottes, 1988; Grantz et al., 1990; Embry, 1990; Plafker and Berg, 1994;

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Figure 5 The Arctic Ocean with the sense of rotation and pole of rotation indicated for the rotational plate tectonic interpretation.

Figure 5). Recently, Lane (1997) challenged the validity of this interpretation, pointing out some of its inconsistencies and presented a brief alternative interpretation with multi-phase rifting (Figure 6). Both the rotation and multi-phase rifting interpretations accept that the margins of northern Alaska and the Arctic Islands are rifted. In the rotation interpretation the two margins are conjugate and opened in the Late Hauterivian (Cretaceous), with some continuation of extensional faulting into the Albian along the Tuktoyaktuk Peninsula. In Lane’s (1994, 1997) multi-phase rifting interpretation there are three main phases of development each with differing spreading centers and poles of rotation. Upper Cretaceous to Tertiary sediments accumulated on the thermally subsiding margin of Canada Basin. However, Late Cretaceous and Tertiary orogenesis encroached from the south and west, producing a thinskinned, low-taper foreland belt extending outward across the basin (Lane, 1998). The combination of thermal subsidence, and sediment and tectonic loading from the orogenic activity has produced a basin with between 14 and 16 km of Upper Cretaceous to Recent sediment.

3. Stratigraphy and Sedimentology 3.1. Stratigraphy Initial stratigraphic studies of the Upper Cretaceous to Tertiary succession in the Beaufort-Mackenzie Basin began on the onshore part of the basin, where a series of alternating shale-rich and sandstone-rich successions allowed a conventional lithostratigraphic scheme to be proposed (Young et al., 1976; Young and McNeil, 1984). However, as more data from the offshore areas became available it was apparent that basin-scale stratigraphic analysis required a different scheme, and a sequence analysis approach was adopted by Dixon et al. (1992). Using reflection seismic and well data they subdivided the Upper Cretaceous to Recent succession into 12 basin-scale,

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Figure 6 Lane’s (1997) interpretation of a multiphase opening of the Arctic Ocean. Stage1: Late Jurassic to Late Hauterivian (Early Cretaceous); Stage 2: Cenomanian (Late Cretaceous); and Stage 3: Late Cretaceous.

transgressive-regressive sequences (Figure 7), each bounded by major unconformities, or surfaces of maximum regression basinward of the unconformities. Dating of the sequences was based primarily on the identification of foraminiferal interval zones (McNeil, 1989, and in Dixon, 1995; McNeil et al., 1990), with additional supporting data using pollens, spores, and dinoflagellates (McIntyre, in Dixon, 1995). Jurassic and Cretaceous strata are identified using conventional lithostratigraphic terminology (Dixon, 1982, 1992a). A number of regionally extensive unconformities have been recognized in the Cretaceous succession and large-scale, widespread, transgressive-regressive sequences have been identified (Dixon, 1993, 1999).

3.2. Sedimentology Jurassic to Aptian strata are characterized by a predominance of shoreface to marine shelf deposits with some local non-marine, deltaic beds in Valanginian strata and fan-delta deposits in Hauterivian to Aptian beds (Dixon, 1992a, Figure 8). During this phase of sedimentation extensional faulting on a continental margin was the prevalent tectonic style. In the Albian there was a shift in the sedimentological regime (Figure 9) following a

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Figure 7 Stratigraphic nomenclature and biostratigraphic zonation scheme (modi¢ed from Dixon et al., 1992; McNeil, 1989, 1997; Harrison et al., 1999). Abbreviations for some sequences and major unconformities used on other ¢gures are indicated.

major transgression onto the craton, which extended far to the south, into western Canada and the U.S.A. In the northern Yukon and under parts of Tuktoyaktuk Peninsula, deep-water troughs developed due to a combination of extensional and compressional tectonics. Into these troughs a variety of coarse- to fine-grained siliciclastic sediments were deposited on submarine fans and basin plains. South and southeast of these troughs was an extensive shelf on which predominantly muds were deposited, with sandstone deposition close to the rising Cordilleran Orogen (Dixon, 1992a, 1992b, 1999). This change in the sedimentary regime during the Albian reflects the earliest influence of the Cordilleran Orogen on sedimentation in northern Canada. Upper Cretaceous to Tertiary deposits in the Beaufort-Mackenzie area range from organic-rich shales deposited on anoxic to dysaerobic outer shelf, slope and possibly basinal environments, non-anoxic slope and basin shales, deep-water submarine fan deposits, shelf shales and sandstones, to deltaic deposits of interbedded sandstone, shale and coal (Dixon et al., 1992; Dixon, 1995). Within these broad depositional realms more detailed paleoenvironments can be readily identified, especially in the deltaic deposits. In the latter case, upper delta-plain, lower delta-plain, and delta-front deposits are recognized. Delta-plain deposits can be further subdivided into distributary channel, crevasse channel, crevasse splay, and overbank deposits. Within the transgressive-regressive sequences, individual prograding delta successions can be recognized from their gammaray log signature (Figure 10), each forming an overall coarsening-upward succession of prodelta shale, gradationally overlain by delta-front sandstone, in turn overlain by delta-plain deposits, usually erosionally overlying the delta-front beds. This type of succession is repeated throughout the deltaic deposits of the sequences. Sandstones of lower delta-plain to delta-front origin have proven to be excellent reservoirs for oil and gas in several sequences, especially where they are overlain by thick prodelta shales (e.g., the Taglu gas field, Dixon, 1981). Shelf and prodelta deposits are dominated by silty to sandy shales with generally very thin interbeds of sandstone. At the distal ends of submarine delta lobes, thin to medium beds of sandstone beds may be deposited on the shelf/prodelta environment, and are locally hydrocarbon bearing (e.g., Tarsiut). Submarine fan deposits are known from at least two of the sequences associated with the Fish River and Kugmallit sequences. Seismic interpretation of data from the west Beaufort suggest that some Taglu sequence sandstone beds in the Edlok N-56 well may be submarine fan deposits (Dietrich et al., 1989a). Submarine fan beds of the Fish River (Dixon, 1988;

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Late Berriasian (Lower Cretaceous) depositional environments (from Dixon, 1992a, Figure 54).

Myers, 1994) and Taglu sequences are generally thin and sandy, whereas those of the Kugmallit sequence are thick and muddy (Dixon et al., 1984). Several of the sequences have been extensively eroded and either have had their deltaic deposits removed (e.g., Akpak sequence), or have an areally restricted and limited amount of preserved deltaic strata (e.g., Richards and Mackenzie Bay sequences). The three oldest sequences, Boundary Creek, Smoking Hills, and Mason River, are outer shelf, slope and basinal deposits, with their equivalent deltaic beds preserved far to the south, on the craton (Dixon, 1992a, 1992b). The Boundary Creek and Smoking Hills beds contain organic-rich shales, deposited in low-oxygen slope and basinal environments.

3.3. Stratigraphic history Lower Jurassic to Aptian strata are characterized by their epicontinental character, with the sediment source primarily from the craton to the southeast and east (Figure 8). Shorelines tend to trend southwest, commonly paralleling positive tectonic elements, such as the Eskimo Lakes Arch (Dixon, 1992a). Quartz-rich sandy sediment was deposited along this shoreline trend, grading westward and northwestward into shelf muds. In the Late Aptian and into the Albian, the influence of Cordilleran tectonics became evident in the style of sedimentation. A major transgression to the southeast and south expanded the basin limits, producing an extensive muddy shelf, south and southeast of Mackenzie Delta. In the northern Yukon the effects of tectonic loading from orogenic activity in adjacent Alaska produced curvilinear deep-water troughs which were filled with sediment-gravity-flow and pelagic deposits (Figure 9). The Albian period marks a transition between the Jurassic to Aptian, rift-dominated phase and the post-Albian compression-dominant phase. Although the main phase of sedimentation in the Beaufort-Mackenzie Basin did not begin until the Late Maastrichtian, there was a preceding phase of Cenomanian to Early Maastrichtian sedimentation during which a

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Early Albian depositional environments (from Dixon, 1992a, Figure 55).

thin, irregularly developed veneer of outer shelf to slope, possibly basinal, commonly organic-rich, muds were deposited on the continental margin (Boundary Creek, Smoking Hills, and Mason River sequences). These represent the distal deposits of a major influx of sediment derived from the orogen to the south. The locus of sedimentation during the Cenomanian to Middle Maastrichtian was on the craton, in a broad foreland basin — the Porcupine-Mackenzie Basin (Dixon, 1992a, 1992b, 1999; Dixon et al., 1992). A major northward shift in sedimentation in the Late Maastrichtian was preceded by the development of an unconformity, indicating the end of a pulse of tectonism within the evolving Cordillera. From the Late Maastrichtian until the Pliocene, deposition of large delta complexes dominated sedimentation in the Beaufort-Mackenzie Basin. These delta complexes formed a series of large-scale transgressive-regressive sequences that, through time, migrated oceanward from the southwest margin of the basin in the Early Tertiary to the central part of the basin in the Pliocene (Figures 11–13). Upper Maastrichtian deposits are represented by sandstones and conglomerates of the Cuesta Creek Member (the member also represents the basal beds of the Fish River sequence) which were deposited on a submarine fan (Dixon, 1988; Myers, 1994). As sediments prograded northward, the Upper Maastrichtian submarine fan deposits were succeeded by slope and shelf muds of the lower Fish River sequence. These in turn are overlain by a thick succession of fluviodeltaic strata of Early Paleocene age (upper part of the Fish River sequence). The Fish River sequence was the first of the major deltaic complexes to be deposited in the Beaufort-Mackenzie Basin (Figures 12 and 13a). Subsequent deposition can be divided into three main tectono-stratigraphic phases, each phase bounded by a major unconformity that marks the culmination of a period of deformation. The first phase ended in the Middle Eocene with the development of an unconformity between the Taglu and Richards sequences. During this Maastrichtian to Mid-Eocene phase, the deltaic complexes were deposited along the southwest basin margin, with the locus of sedimentation for each successive sequence shifting southeastward to eastward, and slightly basinward with time (the Fish River, Aklak and Taglu sequences, Figures 12 and 13). One unusual aspect of this period of deposition is the abundance of well-developed coals in the delta-plain deposits of the Aklak sequence (Late Paleocene to Early Eocene). Although coals are present in other sequences, none compare in quality, thickness, and lateral extent to those in the Aklak sequence. Presumably climatic conditions, vegetation growth, and depositional conditions were optimum for coal formation during the Late Paleocene and Early Eocene (see Moran et al., 2006).

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Figure 10 Typical gamma-ray response and interpreted depositional environments (based on core interpretation) of prograding delta deposits in the Taglu sequence of the Taglu C-42 well (after Dixon, 1981).

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Figure 11 Cross-section illustrating the basinward progradation of delta-complexes in the Beaufort-Mackenzie Basin.

Figure 12 Maximum progradation of delta-front deposits through time. (1) Fish River sequence (Paleocene), (2) Aklak sequence (Early Eocene), (3) Taglu sequence (Middle Eocene), Richards sequence (Late Eocene), (4) Kugmallit sequence (Late Oligocene), and (5) Iperk sequence (Pliocene).

The second tectono-stratigraphic phase lasted until the Late Miocene, during which the Richards, Kugmallit, Mackenzie Bay, and Akpak sequences were deposited. During this period the locus of sedimentation shifted basinward from its initial location in the south-central part of the basin (Figures 12 and 13c). In Late Eocene time a major base-level drop led to extensive erosion of the Richards sequence, removing most of the Richards deltaic deposits. This base-level fall also exposed the Late Eocene shelf to subaerial erosion, resulting in deposition directly into deep-water and the development of a major submarine fan during the latest Eocene or earliest

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Figure 13 Distribution of preserved depositional facies of the: (a) Fish River, (b) Taglu, and (c) Kugmallit sequences at maximum progradation.

Oligocene. Subsequent base-level rise and progradation resulted in the development of one of the most extensive deltaic complexes in the basin, the Kugmallit sequence (Figure 13c), which is up to 4,000 m thick and occupies a large area under the central Beaufort Shelf. Miocene sedimentation (Mackenzie Bay and Akpak sequences) saw a drastic reduction in the volume of sediment and the development of less extensive delta complexes. At the end of the Miocene a major erosional unconformity developed, marking the culmination of another tectonic episode. During the Pliocene a very thick succession of prograding sediments created the modern shelf geometry (Iperk sequence, Figures 11 and 12). This tectono-stratigraphic phase is characterized by the lack of

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deformation, although in contiguous offshore Alaska deformation continued, affecting strata as young as the Recent. Also, this phase of sedimentation saw a significant basinward shift in the locus of deposition (Figure 12). The Late Miocene unconformity also marks a major change in climate between the Miocene and Pliocene, when the climate cooled significantly. The pattern of delta migration (Figure 12), distribution of depositional facies within each sequence (Figure 13), litharenite composition of sandstones, and abundance of coarse-grained siliciclastics within the Tertiary deltaic beds indicate that the main source of sediment was to the southwest, from the northward migrating Cordilleran Orogen. This contrasts with the modern Mackenzie River system which has its origins on the Canadian Shield to the southeast and extends across the Northern Interior Plains, with tributaries from the mountain belt. The modern Mackenzie Delta formed subsequent to the last glaciation. The resulting sea-level rise and transgression after the last glaciation shifted sedimentation southward. Also, the modern drainage system of the Mackenzie River is a post-glacial phenomena (Duk-Rodkin and Hughes, 1994), during which a very long river system formed. A result of having such a long drainage system is that many of the coarse-grained siliciclastic components are deposited in lakes and low-gradient parts of the river. Thus the modern Mackenzie Delta is dominated by clay to very fine-grained sand, in contrast to the Tertiary deltas, in which a wide range of grain sizes were deposited.

4. Structure and Tectonics 4.1. Basin-margin faults On the southeast margin of the basin are two prominent fault zones, the Eskimo Lakes Fault Zone (ELFZ) (Figures 3 and 14) and the Outer Hinge Line (OHL) (Figure 3). These are predominantly pre-mid-Cretaceous fault systems related to rifting and continental breakup. However, they underwent some reactivation during the Late Cretaceous and Tertiary. Between these two fault zones is the Kugmallit Trough, a pre-mid-Cretaceous halfgraben that also acted as a slight depression during Late Cretaceous to Tertiary deposition. These two fault zones merge to the northeast, offshore from Banks Island (Dixon et al., 1992, Figure 3) to form a single hingeline. Deep-reflection seismic profiles show the faults of the ELFZ to have a listric character, soling deep within Precambrian rocks (Cook et al., 1987). The nature of the faults at crustal depths in the OHL are not as well imaged as those of the ELFZ, but the two zones merge into a single, narrow transition zone 50–60 km wide along the basin’s southeastern margin (Dietrich et al., 1989b). Modeling of gravity and aeromagnetic data, tied to regional deep-crustal seismic reflection and refraction data, demonstrate that the zone of transition between continental and oceanic crust is narrow and abrupt in the southeast, but is much broader in the west (Stephenson et al., 1994).

4.2. Beaufort Foldbelt The prominent folds in the Beaufort Foldbelt vary from linear to curvilinear in plan view, with strike lengths of up to 130 km and wavelengths from 10 to 20 km (Figure 3). In the Mackenzie Delta area anticlines plunge generally to the northwest but in the western Beaufort Sea they tend to plunge to the northeast. Anticlines are large amplitude structures with up to 3 km of relief and generally have rounded hinges, although a few chevronstyle hinges occur locally. Most anticlines are asymmetric, with the steeper limbs commonly facing basinward and many are cored with north- or northeast-verging thrust faults (Figure 15). South- or southeast-verging back thrusts occur locally. Most thrust faults flatten into bedding planes. A number of regionally important de´collement horizons have been identified from surface and reflection seismic studies. These occur primarily in shale units, such as an unnamed Cambrian succession, the Carboniferous Kayak Formation, the Jurassic Kingak Formation, the Upper Cretaceous Boundary Creek and Smoking Hills formations, and in Paleocene and Eocene units (Lane and Dietrich, 1995). In the Beaufort-Mackenzie Basin the most prominent de´collement is within the Upper Cretaceous strata, with lesser detachments in Lower Tertiary strata. Some recent workers have suggested that shale diapirism is associated with the thrust and foldbelt (Bergquist et al., 2003). Within the foldbelt are two broad uplifts, the Herschel and Blow River highs (Figure 3), defined by the antiform geometry of a prominent Mid-Eocene unconformity. Below this unconformity are intensely folded and thrust faulted strata. The Herschel High has an associated synclinal feature, the Demarcation Sub-basin, on its southern, landward, flank which has been interpreted as a piggyback basin (Dietrich and Lane, 1992). The intense deformation seen in the Blow River High continues onshore where it is expressed in the highly deformed Albian shales of the Blow Trough (also called the Rapid Depression by some authors).

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Figure 14 Re£ection seismic pro¢le illustrating normal faults in the Eskimo Lakes Fault Zone on Tuktoyaktuk Peninsula. Faults o¡set Paleozoic (P) and Jurassic--Lower Cretaceous (Jkl) strata, and at deeper levels the fault planes parallel basement (Proterozoic) re£ections, i.e., become listric.

Figure 15 Re£ection seismic pro¢le illustrating asymmetric, thrust-cored anticlines in Lower Tertiary strata, western Beaufort Sea. Identi¢ed unconformities are: Middle Eocene, base-Richards sequence (mE), Upper Eocene, base-Kugmallit sequence (uE), and Upper Miocene, base-Iperk sequence (uM). Syntectonic wedges (sw) of the Eocene Taglu sequence are apparent on the southwestern (landward) £anks of the anticlines.

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4.3. Listric faults The central part of the basin is characterized by the presence of listric faults, most of which face basinward and sole out in Upper Cretaceous and Lower Tertiary strata (Figure 16). These faults have throws of up to 1,500 m on the larger ones and trend east to northeast, at a high angle to the underlying fold axes. Associated with the listric faults are rotated fault blocks and antithetic faults. Many show multiple phases of growth. Because contractional deformation and listric faulting were commonly contiguous, some faults show reverse displacement on an original normal displacement, resulting in ‘‘pop-up’’ structures. Many of the larger faults occur in two prominent fault zones, the TFZ and TAFZ (Figures 3 and 4b). The TFZ appears to merge northeastward into the OHL and is located on the north side of the Tununuk High. These two latter features are related to pre-Upper Cretaceous rift development of the basin margin which is, in turn, interpreted as controlling the location of the Eocene listric faults in the TFZ. The TFZ separates the young, weak, thick basin-fill succession from the older, more rigid basin-margin strata to the southeast (Lane and Dietrich, 1995). East-trending listric faults merge onto the main TFZ (Figure 3). Farther north, the east end of the TAFZ abuts the OHL at a high angle and there is no obvious genetic link, as is seen with the TFZ, other than to act as a buttress against which faults in the TAFZ terminated. The TAFZ is characterized by converging and diverging fault strands producing wedge- or lens-shaped map patterns (Figure 3). Both the TFZ and TAFZ terminate at their western ends against the Blow River High. The TAFZ is interpreted to have resulted from Oligocene to Miocene deformation interacting with a complex array of boundary conditions, including lithospheric heterogeneity, proximity to the basin margin and also to the basin depocentre (Stephenson et al., 1994; Lane and Dietrich, 1995).

4.4. Timing of deformation Based on the structural and stratigraphic interpretation of reflection seismic data tied to about 260 wells, many with detailed biostratigraphic data, five major unconformities define the culmination of five phases of significant deformation, these are Early, Middle, and Late Eocene, Late Oligocene, and Late Miocene, corresponding to the base of the Taglu, Richards, Kugmallit, Mackenzie Bay, and Iperk sequences, respectively. Although these unconformities marked the end of major tectonic pulses there has been a general continuum of deformation in

Figure 16 Re£ection seismic pro¢le illustrating listric-growth faults in Tertiary strata, eastern Beaufort Sea. Identi¢ed unconformities are: Middle Cretaceous (mC), Upper Eocene, base-Kugmallit sequence (uE), Upper Oligocene, base-Mackenzie Bay sequence (uO), and Upper Miocene, base-Iperk sequence (uM). Listric faults sole into sub-horizontal detachments in Lower Tertiary or Upper Cretaceous strata, at re£ection times of about 6.0 s (about 10--12 km).

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the basin, as seen in the stratal onlap and convergence on to many structures at various stratigraphic levels and in the growth of many listric faults (Figures 15 and 16). The oldest pulse of deformation is best seen in the western part of the basin where Upper Paleocene to Lower Eocene Aklak strata onlap the flanks of folds and uplifts. The overlying unconformity at the base of the Taglu sequence marks the culmination of this phase of deformation (Early Eocene). The Middle Eocene unconformity (base of the Richards sequence) is widespread and marks the end of the most intense phase of deformation in the basin. In the southwestern part of the basin many structures are deeply truncated by this unconformity. Above the unconformity younger strata are less severely deformed. Deformation continued from the Late Eocene to Miocene with deformation extending into progressively younger strata basinward. The listric faults of the TFZ appear to have developed mostly from the Middle Eocene to the Late Eocene, as is indicated by the thickening of Taglu and Richards strata across the faults. Late Miocene deformation within the TFZ resulted in some faults inverting due to compression, an important clue to the tectonic evolution (Lane and Dietrich, 1995). A Late Eocene unconformity was accompanied by significant base-level fall and the subsequent deposition of extensive submarine fan deposits. Faulting in the TAFZ post-dated Eocene deformation, with most of its growth occurring during the Oligocene and Miocene. Only minor reactivation of the TAFZ occurred during the Pliocene. The youngest phase of deformation culminated in the Late Miocene with the development of a major, basinwide unconformity (base of Iperk sequence) that resulted in large volumes of sediment being eroded and the development of an extensive pediment surface. Deformation in the Pliocene to Recent has been minimal in the Beaufort-Mackenzie Basin, although in contiguous offshore Alaska significant deformation continued, with some far offshore structures extending to the seafloor (Dinter et al., 1990; Grantz et al., 1990). Fission track data from land-based samples in northern Yukon are dominated by Early Eocene cooling ages (O’Sullivan and Lane, 1997), whereas in adjacent Alaska, which is on strike with the offshore Beaufort Foldbelt, similar studies gave cooling ages of Late Paleocene, Early and Middle Eocene, and Early Miocene (O’Sullivan et al., 1993), most of which correspond to events recorded in the basin stratigraphy and structural development, except for the Early Miocene event which appears to have no corresponding stratigraphic or structural signature in the basin. The Middle Eocene and Early Miocene cooling events have been recorded from a number of stations, suggesting regional tectonic events. Overpressured strata are widespread in the basin (Hitchon et al., 1990) and appear to have existed since the beginning of basin development and is most likely due to the high rate of sedimentation and rapid subsidence (Issler, 1992). Overpressuring probably aided deformation, allowing strata to deform at lower stress levels and to favor brittle deformation (i.e., thrust faulting), and the development of through-going de´collements (Lane and Dietrich, 1995). Also, overpressured shales may have resulted in shale diapirism.

5. Economic Geology Oil and gas are the most valuable resource assets of the basin, with 53 discoveries to-date. Coal is also common in some sequences, especially the Upper Paleocene to Lower Eocene Aklak sequence, but it is not considered a viable economic resource due to the difficulty of access — most of the coal resources being offshore at considerable subsea depths. A limited amount of coal from the outcropping Aklak sequence was mined for local use from a surface adit along the western margin of Mackenzie Delta during the 1950s. When dieselpowered electrical generators and heating oil became more available to the communities in the 1960s the mine was abandoned. Oil and gas exploration began in the 1960s with surface studies and acquisition of reflection seismic. The first well was drilled in 1965 on Mackenzie Delta and drilling continued into the early-1990s on a sporadic basis, with the peak of drilling in the 1970s. Very little exploration took place during the 1990s, although in the last year of that decade a small gas field about 50 km north of Inuvik was put on production for local consumption (Ikhil N-35, Figure 2). In late-1999 a new phase of increased exploration began with the issuance of new exploration permits. The first discovery was an oil well, the Atkinson Point H-25 (Figure 2, Morrell, in Dixon,1995), drilled on central Tuktoyaktuk Peninsula in 1969. Oil was recovered from Lower Cretaceous sandstones and conglomerates of the Atkinson Point Formation. This initial discovery led to a spurt of exploration and the subsequent discovery of a number of oil and gas fields on Mackenzie Delta, some of which have substantial gas reserves (e.g., Parsons Lake with 51.7 109 m3 and Taglu with 85 109 m3, Figure 2). In the mid to late 1970s exploration moved into the shallow offshore waters north of the delta, initially using various types of man-made islands as drilling platforms but resorting to ice-strengthened drillships in deeper

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water. Drilling in water depths where drillships need to be used is limited to about a three to four month season of open water (approximately August to November). A number of oil and gas discoveries were made in the deeper waters (e.g., Kopanoar, Nerlerk, Koakoak, Kenalooak, Figure 2). However, ice conditions can be variable from year to year and some areas can remain icebound in particularly bad years. Drilling activity declined during the late 1980s and stopped for several years from about 1990, by which time about 250 wells had been drilled. Three broad groups of structural traps contain the bulk of the discovered hydrocarbons: 1. Structures associated with the rifted margins where the reservoirs are usually in Lower Cretaceous sandstones (e.g., Atkinson Point, Parsons Lake), or fractured Paleozoic carbonates (e.g., Mayogiak) (Figure 17a). 2. Rotated fault blocks associated with listric faults (e.g., Taglu, Amauligak, Unipkat, Isserk) (Figure 17b). 3. Faulted anticlines of the Beaufort Foldbelt (e.g., Adlartok, Kopanoar, Koakoak) (Figure 17c). In the second and third category the reservoirs are in Tertiary sandstones, most commonly in the Taglu and Kugmallit sequences, and less frequently in the Aklak sequence. Tertiary reservoirs are predominantly in deltafront and lower delta-plain sandstones and some of the most prolific reservoirs are where such sandstones are overlain by a thick accumulation of prodelta shales, which acts as an efficient top seal (e.g., Taglu and Amauligak). Some discoveries have been in shelf (or distal delta-front, e.g., Tarsiut) and deep-water sandstones (e.g., Kopanoar), but these reservoirs are less prolific and less continuous than delta-front/delta-plain sandstones. Untested structural traps include shale diapirs (Figure 17d). Sandstone reservoirs vary in quality with age, grain size, and depositional environment. Lower Cretaceous sandstones are predominantly quartz arenites and well cemented (usually quartz and/or calcite), consequently porosity and permeability values tend to be low. However, where secondary porosity has developed porosity can be between 15 and 20%, such as in the Parson Lake gas field. Most of the known Cretaceous sandstones are

Figure 17

Schematic representation of various structural traps in the Beaufort-Mackenzie Basin.

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shoreline to inner shelf deposits, with local terrestrial beds. Tertiary sandstones are mostly litharenites whose main components are quartz and chert with variable amounts of metamorphic, sedimentary, and volcanic rock fragments (Young, 1975; Schmidt, 1987; Nentwich and Yole, 1982). Finer grained sandstones, especially those deposited in shelf and deep-water environments may contain a clay matrix. The dominant cements include calcite, dolomite, siderite, and kaolinite. Porosity and permeability values are generally high in Tertiary sandstones. Porosity values of 12–25% are common, and up to 30% is not uncommon in some Oligocene sandstones. Sandstones in the Aklak and Taglu sequences (Upper Paleocene to Lower Eocene) tend to be, on average, less porous than those of the Oligocene Kugmallit sequence. Also, finer grained sandstones generally are less porous than coarser ones. As previously indicated, delta-front and lower delta-plain sandstones are the best reservoirs. Much of the high porosity in the Tertiary sandstones is due to the dissolution of cements, especially carbonate cements. Source rocks within the basin include the organic-rich, Upper Cretaceous Boundary Creek and Smoking Hills sequences (Snowdon, 1978, 1980a), which can contain up to 12% organic matter. Some of the oil discoveries in Cretaceous strata along Tuktoyaktuk Peninsula have been derived from these rocks (Powell and Snowdon, 1975; Snowdon and Powell, 1979). Geochemical screening of potential source rocks within the Tertiary succession initially failed to identify a source rock because most of the shale successions are lean in organic matter and tend to be gas-prone. It was not until the technique of identifying geochemical marker compounds was employed that a possible Tertiary source rock was identified. The basal beds of the Richards sequence were found to contain a marker compound that is also present in many of the Tertiary oils (Brooks, 1986a, 1986b). The marker compound present in the basal Richards sequence had not been noted in other successions but subsequent to Brooks’ (1986a, 1986b) work it has been identified in the Taglu sequence (Snowdon et al., 2004). An alternative to the conventional hydrocarbon-generation model was proposed by Snowdon (1980b) and Snowdon and Powell (1982), who suggested that the occurrence of significant amounts of resinite in the sediments would allow hydrocarbon generation under the conditions found in the basin. Recent work on the geochemical characterization of the oils suggests that they may be derived from the known Upper Cretaceous source rocks but have been modified by subsequent migration through the Tertiary succession (Li et al., 2006). The low thermal maturity of the basin, the generally low organic content, and the predominance of Type III terrestrial organic matter would not normally be good indicators for significant hydrocarbon generation yet the Beaufort-Mackenzie Basin contains large volumes of oil and gas. Kinetic modeling by Issler and Snowdon (1990) suggests that the low thermal gradient only extends down to about 5 km, below which the thermal gradient rises sharply. Their model predicts that with only a 1–1.5 km of additional burial the sediment column would pass into and through the oil generation window. These conditions were interpreted to be the result of rapid deposition of the Pliocene Iperk sequence and subsequent thermal disequilibrium. The modeling also suggests that a modest amount of oil generation occurred at an early stage of burial of the Richards sequence but that the bulk of the oil generation occurred after the Iperk sequence was deposited. With the exception of the small gas field at Ikhil, north of the town of Inuvik, none of the discoveries are producing and only a few of the larger discoveries have been delineated with step-out wells. The three largest discoveries are the Parsons Lake gas field (Lower Cretaceous sandstone reservoirs), Taglu gas field (Middle Eocene sandstone reservoirs), and the Amauligak oil and gas field (Oligocene sandstone reservoirs) (see Morrell, in Dixon, 1995). The resources in the Beaufort-Mackenzie Basin undoubtedly will be placed on-stream as the hydrocarbons from the more accessible parts of western Canada become depleted and cannot meet the demand for oil and gas.

6. Summary The Beaufort-Mackenzie Basin is a Late Cretaceous to Recent, post-rift, continental-margin basin that has been overprinted by compressional tectonics associated with the formation of the adjacent Cordillera. Although the post-rift stage began in the Cenomanian the first sediments were a thin blanket of muds that continued to be deposited into the Middle Maastrichtian. The first major pulse of sedimentation on the continental margin did not begin until the Late Maastrichtian. Subsequent to the Late Maastrichtian at least seven delta complexes were deposited in a series of large-scale transgressive-regressive sequences which, over time, prograded basinward, finally creating the modern Beaufort Shelf. Contemporaneous with sedimentation was the development of compressional folds and thrust faults, and extensional listric faults. Five major pulses of deformation have been identified, but two well-developed unconformities (Mid-Eocene and Late Miocene) divide the Upper Cretaceous to Recent strata into three major tectono-stratigraphic assemblages. The Mid-Eocene unconformity marks the end of the most intense period of deformation within the basin. The Late Miocene unconformity

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marks the end of intense deformation throughout most of the basin and strata above this unconformity are generally undeformed to only mildly deformed. The Upper Cretaceous and Tertiary basin-fill and pre-Cenomanian rifted strata contain large volumes of oil and gas, and significant coal resources. Large anticlines, fault-related structures and porous sandstones combine to produce good hydrocarbon traps. Upper Cretaceous strata contain some very organic-rich intervals but the Tertiary sediments generally are low in organic carbon. Thermal maturity is generally low to the depths drilled to-date. Recent geochemical studies indicate that the oils may have been derived initially from deeply buried Upper Cretaceous source rocks but that subsequent migration through the Tertiary succession has modified their geochemical signature.

ACKNOWLEDGMENTS This summary of work on the Beaufort-Mackenzie Basin is the result of many years of study by the authors and we wish to thank the Geological Survey of Canada, especially the staff at the Calgary office, for their support and constructive criticism. For this particular manuscript we appreciate the comments by Ashton Embry, William Galloway, and Andrew Miall which helped to clarify several issues. Earth Sciences Sector contribution number 20060510.

REFERENCES Bergquist, C. L., Graham, P. P., Johnston, D. H., and Rawlinson, K. R., 2003, Canada’s Mackenzie Delta: fresh outlook at an emerging basin. Oil and Gas Journal, v. 101(November), pp. 42–46. Brooks, P. W., 1986a, Biological marker geochemistry of oils from the Beaufort-Mackenzie region, Arctic Canada. Bulletin of Canadian Petroleum Geology, Calgary, Canada, v. 34, pp. 490–505. Brooks, P. W., 1986b, Unusual biological marker geochemistry of oils and possible source rocks, offshore Beaufort-Mackenzie Delta, Canada. Organic Geochemistry, Elsevier, Amsterdam, v. 10, pp. 401–406. Chamney, T. P., 1969, Abnormally thick Tertiary-Cretaceous sequence, Mackenzie Delta, District of Mackenzie, in Current Research, Part B, Geological Survey of Canada, Ottawa, Canada, Paper 69-1B, pp. 69–72. Chamney, T. P., 1973, Tuktoyaktuk Peninsula Tertiary and Mesozoic biostratigraphy correlations, in Report of Activities, Part B; Geological Survey of Canada, Ottawa, Canada, Paper 73-1B, pp. 171–178. Cook, F. A., Coflin, K. C., Lane, L. S., Dietrich, J. R., and Dixon, J., 1987, Structure of the southeast margin of the Beaufort-Mackenzie Basin, Arctic Canada, from crustal seismic reflection data. Geology, Boulder, U.S.A., v. 15, pp. 931–935. Dietrich, J. R., Dixon, J., McNeil, D. H., McIntyre, D. J., Snowdon, L. R., and Cameron, A. R., 1989a, Geology, biostratigraphy and organic geochemistry of the Natsek E-56 and Edlok N-56 wells, Beaufort Sea, Arctic Canada. In Current Research, Part G; Geological Survey of Canada, Ottawa, Canada, Paper 89-1G, pp. 133–157. Dietrich, J. R., Coflin, K. C., Lane, L. S., Dixon, J., and Cook, F. A., 1989b, Interpretation of deep seismic reflection data, Beaufort Sea, Arctic Canada. Geological Survey of Canada Open File 2106, Ottawa, Canada, 15 pp. Dietrich, J. R., and Lane, L. S., 1992, Geology and structural evolution of the Demarcation Subbasin and Herschel High, western Beaufort-Mackenzie Basin, Arctic Canada. Bulletin of Canadian Petroleum Geology, Calgary, Canada, v. 40, pp. 188–197. Dinter, D. A., Carter, L. D., and Brigham-Gettes, J., 1990, Late Cenozoic evolution of the Alaskan north slope and adjacent continental shelves, in Grantz, A., Johnson, L., and Sweeeny, J. F. eds., The Arctic Ocean Region, Geological Society of America, The Geology of North America, Boulder, U.S.A., v. L, pp. 459–490. Dixon, J., 1981, Sedimentology of the Eocene Taglu Delta, Beaufort-Mackenzie Basin: example of a river–dominant delta, Geological Survey of Canada, Ottawa, Canada, Paper 80-11, 11 pp. Dixon, J., 1982, Subsurface stratigraphy of the Mackenzie Delta-Tuktoyaktuk Peninsula, N.W.T. Geological Survey of Canada, Ottawa, Canada, Bulletin 349, 52 pp. Dixon, J., 1988, Depositional setting of the Maastrichtian Cuesta Creek Member, Tent Island Formation, northern Yukon, in Current Research, Part D, Geological Survey of Canada, Ottawa, Canada, Paper 88-1D, pp. 61–65. Dixon, J., 1992a, A review of Cretaceous and Tertiary stratigraphy in the northern Yukon and adjacent Northwest Territories, Geological Survey of Canada, Ottawa, Canada, Paper 92-9. Dixon, J., 1992b, Mesozoic stratigraphy, Eagle Plain area, northern Yukon, Geological Survey of Canada, Bulletin 408, Ottawa, Canada. Dixon, J., 1993, Regional unconformities in the Cretaceous of northwest Canada. Cretaceous Research, v. 14, pp. 17–38. Dixon, J., 1995, ed., Geological atlas of the Beaufort-Mackenzie area, Geological Survey of Canada, Miscellaneous Report 59, Ottawa, Canada, 173 pp. Dixon, J., 1999, Mesozoic-Cenozoic stratigraphy of the Northern Interior plains and plateaux, Northwest Territories, Geological Survey of Canada, Ottawa, Canada, Bulletin 536, 56 pp. Dixon, J. D., McNeil, D. H., Dietrich, J. R., Bujak, J. P., and Davies, E. H., 1984, Geology and biostratigraphy of the Dome Gulf et al. Kopanoar M-13 well, Beaufort Sea, Geological Survey of Canada, Ottawa, Canada, Paper 82-13, 28 pp. Dixon, J., Dietrich, J. R., and McNeil, D. H., 1992, Upper Cretaceous to Holocene sequence stratigraphy of the Beaufort-Mackenzie and Banks Island areas, northwest Canada, Geological Survey of Canada, Ottawa, Canada, Bulletin 407, 90 pp. Dixon, J., Morrell, G. R., Dietrich, J. R., Taylor, G. C., Procter, R. M., Conn, R. F., Dallaire, S. M., and Christie, J. A., 1994, Petroleum resources of the Mackenzie Delta and Beaufort Sea, Geological Survey of Canada, Ottawa, Canada, Bulletin 474, 52 pp. Duk-Rodkin, A., and Hughes, O. L., 1994, Tertiary-Quaternary drainage of the pre-glacial Mackenzie Basin. Quaternary International, v. 20, pp. 221–241.

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Embry, A. F., 1990, Geological and geophysical evidence in support of the hypothesis of anticlockwise rotation of northern Alaska. Marine Geology, v. 93, pp. 317–329. Embry, A. F., and Dixon, J., 1990, The breakup unconformity of the Amerasia Basin, Arctic Ocean: evidence from Arctic Canada. Geological Society of America Bulletin, v. 12, pp. 1526–1534. Grantz, A., May, S. D., and Hart, P. E., 1990, Geology of the Arctic continental margin of Alaska, in Grantz, A., Johnson, L., and Sweeney, J. F. eds., The Arctic Ocean region, Geological Society of America, Geology of North America, Boulder, U.S.A., v. L, pp. 257–258. Harrison, J. C., Mayr, U., McNeill, D. H., Sweet, A. R., McIntytre, D. J., Eberle, J. J., Harington, C. R., Chalmers, J. A., Dam, G., and Nohr-Hansen, H., 1999, Correlation of Cenozoic sequences of the Canadian Arctic region and Greenland: implications for the tectonic history of northern North America. Bulletin of Canadian Petroleum Geology, v. 47, pp. 223–254. Hawkings, T. J., and Hatlelid, W. G., 1975, The regional setting of the Taglu gas field, in Yorath, C. J., Parker, E. R., and Glass, D. J. eds., Canada’s Continental Margins, Canadian Society of Petroleum Geologists, Calgary, Canada, Memoir 4, pp. 633–648. Hitchon, B., Sauveplane, C. M., Underschultz, J. R., and Bachu, S., 1990, Hydrogeology, geopressures and hydrocarbon occurrences, Beaufort-Mackenzie Basin. Bulletin of Canadian Petroleum Geology, v. 38, pp. 215–235. Issler, D. R., 1992, A new approach to shale compaction and stratigraphic restoration, Beaufort-Mackenzie Basin and Mackenzie Corridor, northern Canada. American Association of Petroleum Geologists Bulletin, v. 76, pp. 1170–1189. Issler, D. R., and Snowdon, L. R., 1990, Hydrocarbon generation kinetics and thermal modeling, Beaufort-Mackenzie Basin. Bulletin of Canadian Petroleum Geology, v. 38, pp. 1–16. Lane, L. S., 1994, A new plate kinematic model of Canada Basin evolution, in Thurston, D. K. and Fujita, K., eds., Proceedings of the 1992 International Conference on Arctic Margins, U.S. Minerals Management Service OCS Report 94-0040, Anchorage, U.S.A., pp. 283–288. Lane, L. S., 1997, Canada Basin, Arctic Ocean: evidence against a rotational origin. Tectonics, v. 16, pp. 363–387. Lane, L. S., 1998, Latest Cretaceous–Tertiary tectonic evolution of northern Yukon and adjacent Arctic Alaska. American Association of Petroleum Geologists Bulletin, v. 82, pp. 1353–1371. Lane, L. S., and Dietrich, J. R., 1995, Tertiary structural evolution of the Beaufort Sea-Mackenzie Delta region, Arctic Canada. Bulletin of Canadian Petroleum Geology, v. 43, pp. 293–314. Lawver, L. A., and Scotese, C. R., 1990, A review of tectonic models for the evolution of the Canada Basin, in Grantz, A., Johnson, L., and Sweeney, J. F. eds., The Arctic Ocean, Geological Society of America, The Geology of North America, Boulder, U.S.A., v. L, pp. 593–618. Lerand, M. M., 1973, Beaufort Sea, in McCrossan, R. G. ed., The Future Petroleum Provinces of Canada, Canadian Society of Petroleum Geologists, Calgary, Canada, Memoir 1, pp. 315–386. Li, M., Achal, S., Chen, Z., Snowdon, L. R., and Issler, D., 2006, Petroleum systems of the Tertiary-reservoired oils in the BeaufortMackenzie Basin: a new perspective from quantitative whole oil GC/MS and saturate GS/MS/MS data, Canadian Society of Petroleum Geologists-Canadian Society of Exploration Geophysicists-Canadian Well Logging Society, Annual Convention, Program and Abstracts, Calgary, Canada, p. 49 (CD-ROM). McNeil, D. H., 1989. Foraminiferal zonation and biofacies analysis of Cenozoic strata in the Beaufort-Mackenzie Basin of Arctic Canada, in Current Research, Part G; Geological Survey of Canada, Ottawa, Canada, Paper 89-1G, pp. 203–223. McNeil, D. H., 1997. New foraminifera from the Upper Cretaceous and Cenozoic of the Beaufort-Mackenzie Basin of Arctic Canada, Cushman Foundation for Foraminiferal Research (Special Publication No.35), Tulsa, U.S.A., pp. 1–95. McNeil, D. H., Dietrich, J. R., and Dixon, J., 1990, Foraminiferal biostratigraphy and seismic sequences: examples from the Cenozoic of the Beaufort-Mackenzie Basin, Arctic Canada, in Hemleben, C., Kaminski, M. A., Kuhnt, W., and Scott, D. B. eds., Paleoecology, biostratigraphy, paleoceanography and taxonomy of agglutinated Foraminifera, Kluwer Academic Publishers, The Netherlands, pp. 859–882. Moran, K., Backman, J., Brinkhuis, H., Clemens, S. C., Cronin, T., Dickens, G. R., Eynaud, F., Gattacceca, J., Jakobsson, M., Jordan, R. W., Kaminski, M., King, J., Koc, N., Krylov, A., Martinez, N., Matthiessen, J., McInroy, D., Moore, T. C., Onodera, J., Oregan, M., Pa¨like, H., Rea, B., Rio, D., Sakamoto, T., Smith, D. C., Stein, R., St. John, K., Suto, I., Suzuki, N., Takahashi, K., Watanabe, M., Yamamoto, M., Farrell, J., Frank, M., Kubik, P., Jokat, W., and Kristoffersen, Y., 2006, The Cenozoic palaeoenvironment of the Arctic Ocean. Nature, v. 441, pp. 601–605. Mountjoy, E. W., 1967, Upper Cretaceous and Tertiary stratigraphy, northern Yukon and northwestern District of Mackenzie, Geological Survey of Canada, Ottawa, Canada, Paper 66-16, 70 pp. Myers, M. D., 1994. Evolution of Late Cretaceous–Early Tertiary depositional sequences in the Beaufort-Mackenzie Basin, Canada, Unpublished PhD thesis, University of Alaska, Fairbanks, U.S.A., 239 pp. Nentwich, F. W., and Yole, R. W., 1982, Sedimentary petrology and stratigraphic analysis of the subsurface Reindeer Formation (Early Tertiary), Mackenzie Delta-Beaufort Sea area, Canada, in Embry, A. E. and Balkwill, H. R. eds., Arctic Geology and Geophysics, Canadian Society of Petroleum Geologists, Calgary, Canada, Memoir 8, pp. 55–82. Norris, D. K., 1981a, Geology: Herschel Island and Demarcation Point, Yukon Territory, Geological Survey of Canada, Map 1514A, 1:250 000 scale. Norris, D. K., 1981b. Geology: Mackenzie Delta, District of Mackenzie, Geological Survey of Canada, Map 1515A, 1:250 000 scale. Norris, D. K., 1981c, Geology: Blow River and Davidson Mountains, Yukon Territory-District of Mackenzie, Geological Survey of Canada, Map 1516A, 1:250 000 scale. Norris, D. K., 1981d, Geology: Aklavik, District of Mackenzie, Geological Survey of Canada, Map 1517A, 1:250 000 scale. O’Sullivan, P. B., and Lane, L. S., 1997, Early Tertiary thermotectonic history of the northern Yukon and adjacent Northwest Territories, Arctic Canada. Canadian Journal of Earth Sciences, v. 34, pp. 1366–1378. O’Sullivan, P. B., Green, P. F., Bergman, S. C., Decker, J., Duddy, I. R., Gleadow, A. J. W., and Turner, D. L., 1993, Multiple phases of Tertiary uplift and erosion in the Arctic National Wildlife refuge, Alaska, revealed by apatite fission track analysis. American Association of Petroleum Geologists Bulletin, v. 77, pp. 359–385. Plafker, G., and Berg, H. C., 1994, Overview of the geology and tectonic evolution of Alaska, in Plafker, G. and Berg, H. C. eds., The Geology of Alaska, Geological Society of America, The Geology of North America, Boulder, U.S.A., v. G-1, pp. 989–1021. Poulton, T. P., Leskiw, K., and Audretsch, A. P., 1982, Stratigraphy and microfossils of Jurassic Bug Creek Group of northern Richardson Mountains, northern Yukon and adjacent Northwest Territories, Geological Survey of Canada, Ottawa, Canada, Bulletin 325, 130 pp.

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Powell, T. G. and Snowdon, L. R., 1975, Geochemistry of oils and condensates from the Mackenzie Delta, NWT. in Report of Activities, Part C; Geological Survey of Canada, Ottawa, Canada, Paper 75-1C, pp. 41–43. Rowley, D. B., and Lottes, A. L., 1988, Plate kinematic reconstruction of the north Atlantic and Arctic, Late Jurassic to Present. Tectonophysics, v. 155, pp. 73–120. Schmidt, V., 1987, Petrological/diagenetic study of Upper Cretaceous and Tertiary strata, Mackenzie Delta, Geological Survey of Canada, Ottawa, Canada, Open File Report 1534. Snowdon, L. R., 1978, Organic geochemistry of the Upper Cretaceous/Tertiary delta complexes of Beaufort-Mackenzie sedimentary basin, northern Canada, Geological Survey of Canada, Ottawa, Canada, Bulletin 291, 46 pp. Snowdon, L. R., 1980a, Petroleum source potential of the Boundary Creek Formation, Beaufort-Mackenzie Basin. Bulletin of Canadian Petroleum Geology, v. 28, pp. 46–58. Snowdon, L. R., 1980b, Resinite-A potential petroleum source in the Upper Cretaceous/Tertiary of the Beaufort-Mackenzie Basin, in Miall, A. D. ed., Facts and principles of world petroleum occurrence, Canadian Society of Petroleum Geologists, Calgary, Canada, Memoir 6, pp. 509–521. Snowdon, L. R., and Powell, T. G., 1979, Families of crude oils and condensates in the Beaufort-Mackenzie Basin. Bulletin of Canadian Petroleum Geology, v. 27, pp. 139–162. Snowdon, L. R., and Powell, T. G., 1982, Immature oil and condensate-modification of hydrocarbon generation model for terrestrial organic matter. American Association of Petroleum Geologists Bulletin, v. 68, pp. 775–788. Snowdon, L. R., Stasiuk, L. D., Robinson, R., Dixon, J., Dietrich, J. R., and McNeil, J. H., 2004, Organic geochemistry and organic petrography of a potential source rock of Early Eocene age in the Beaufort-Mackenzie Basin. Organic Geochemistry, v. 35, pp. 1039–1052. Stephenson, R. A., Coflin, K. C., Lane, L. S., and Dietrich, J. R., 1994, Crustal Structure and Tectonics of the southeastern Beaufort Sea continental margin. Tectonics, v. 13, pp. 389–400. Wielans, J. B. W., 1992, The pre–Mesozoic stratigraphy and structure of Tuktoyaktuk Peninsula, Geological Survey of Canada, Ottawa, Canada, Paper 90-22, 90 pp. Willumsen, P. S., and Cote, R. P., 1982, Tertiary sedimentation in the southern Beaufort Sea, in Embry, A. E. and Balkwill, H. R. eds., Arctic Geology and Geophysics, Canadian Society of Petroleum Geologists, Calgary, Canada, Memoir 8, pp. 43–53. Yorath, C. J., 1973, Geology of Beaufort-Mackenzie Basin and eastern part of Northern Interior Plains, in Pitcher, M. G. ed., Arctic geology, American Association of Petroleum Geologists, Tulsa, U.S.A., Memoir 19, pp. 41–47. Young, F. G., 1975, Upper Cretaceous stratigraphy, Yukon coastal plain and northwestern Mackenzie Delta, Geological Survey of Canada, Ottawa, Canada, Bulletin 249, 83 pp. Young, F. G. and McNeil, D. H., 1984, Cenozoic stratigraphy of Mackenzie Delta, Northwest Territories, NTS 107B, 107C, Geological Survey of Canada, Ottawa, Canada, Bulletin 336, 63 pp. Young, F. G., Myhr, D. W., and Yorath, C. J., 1976, Geology of the Beaufort-Mackenzie Basin, Geological Survey of Canada, Ottawa, Canada, Paper 76-11, 63 pp.

CHAPTER 17

Postscript: What have We Learned and Where Do We Go from Here? Andrew D. Miall

Contents 1. Basin Analysis in North America 1.1. Basin tectonics 1.2. Stratigraphy and sedimentology 2. North American Geology 2.1. Evolution of the craton 2.2. Sediment fluxes 2.3. Other effects of intraplate stress 3. Remaining Problems Acknowledgments References

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Abstract Many of the developments in basin analysis originated in North America, including ideas about isostasy and about geosynclines. The first comprehensive application of plate-tectonic concepts to interpretations of continental geology took place in Newfoundland, and served as a confirmation of Wilson’s revolutionary idea that the Atlantic Ocean ‘‘closed and then re-opened.’’ Later developments regarding foreland basins and controversies about thick-skinned versus thinskinned tectonism were also resolved from North American data, particularly following the advent of the deep-crustal reflection-seismic methods of COCORP and Lithoprobe. Sequence stratigraphy was also developed initially in North America, stimulated by the recognition of widespread packages of strata separated by unconformities that did not correlate to the standard, largely European, chronostratigraphic framework. The major problems of North American Phanerozoic geology have now been largely solved, but significant issues remain, particularly regarding the local and regional histories of terrane amalgamation in the Appalachian and Cordilleran orogens.

1. Basin Analysis in North America 1.1. Basin tectonics As detailed by Mayo (1985) and Bally (1989), many basic ideas about sedimentary basins, including the concept of the geosyncline, and the theory of isostasy, were American inventions, owing to the work of such individuals as James Hall, Dwight Dana, and Clarence Dutton during the nineteenth century. Kay’s (1951) elaboration of the geosyncline concept was concerned with North American sedimentary basins, and constituted the last comprehensive attempt to document and interpret basins before the advent of modern methods and the emergence of plate tectonics. Ettensohn (Chapter 4) discusses this early history with reference to the Appalachian basin, which may be regarded as the type example of a geosyncline, based on the work of Schuchert (1923). Bally (1989, pp. 397–398) noted the work of Drake et al. (1959) ‘‘who first tried to reconcile modern geophysical-oceanographic observations with the geosynclinal concept’’ and that of Dietz (1963) and Dietz and Holden (1974) who were the first to equate Kay’s ‘‘miogeosyncline’’ with the plate-tectonic concept of an Atlantic-type passive continental margin. All of these researchers and all of the developments to which they refer are North American. In fact, one of the major realizations that has emerged from this historical review is that many of the major developments in basin analysis were made by US and Canadian scientists, or by scientists from Sedimentary Basins of the World, Volume 5 ISSN 1874-5997, DOI 10.1016/S1874-5997(08)00017-8

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elsewhere working in North America as, for example, J. F. Dewey and A. B Watts, did early in their careers. The reason for this probably has much to do with the nature of North American geology, as argued later in this section. Following on Wilson’s classic speculation about the opening and closing of the Atlantic Ocean and the origins of the Appalachian–Caledonian orogen (Wilson, 1966), Mitchell and Reading (1969) made one of the first (not North American) attempts to reinterpret the old tectono-stratigraphic concepts of flysch and molasse in terms of the new plate tectonics. John Bird and John Dewey, in two papers published in 1970, completely revolutionized our understanding of the origins of sedimentary basins (and much of the rest of geology) with reference to the geology of the Appalachian orogen, in particular, that portion of it exposed throughout the island of Newfoundland. Dickinson (1971) made reference to all of this work in his own first pass at relating sedimentary basins to plate tectonics. Drawing on the detailed mapping work of such individuals as Ward Neale, Bob Stephens, Mike Kennedy and Hank Williams, Bird and Dewey (1970), and Dewey and Bird (1970) showed how miogeosynclines, eugeosynclines and the many other types of basin elucidated by Marshall Kay, could be understood with reference to the new, dynamic concepts of plate tectonics. For example, central Newfoundland was interpreted as an arc complex, while western Newfoundland, including the Great Northern Peninsula, were understood to be underlain by Grenville basement, flanked by an Early Paleozoic Atlantic-type extensional continental margin over which had been thrust an allochthonous complex (part of the Taconic orogen) that included significant remnants of oceanic crust. Continental terranes formerly separated by a wide ocean were now in close juxtaposition. Williams (1978) later published his classic map, in which four formal subdivisions of Appalachian geology, first recognized and named in Newfoundland, were extended along the length of the exposed Appalachians as far as Florida. These soon became the basis for comprehensive plate-tectonic interpretations of the Appalachian orogen, which included speculations about the diverse continental sources of continental fragments — the so-called suspect terranes (Williams and Hatcher, 1982; Williams, 1984). Subsequent work, discussed briefly in the concluding section of this chapter, has shown that the Dunnage terrane, as originally defined, is far more complex than at first thought, containing the remnants of several small oceans and two separate terrane now called the Notre Dame terrane and Dashwoods microcontinent, the latter representing a fragment of Grenville crust rifted away from the continent during the breakup of Rodinia (Percival et al., 2004; van Staal, 2005; Zagorevski et al., 2006; Ettensohn, Chapter 4). These breakthroughs of the 1970s initiated a worldwide explosion of studies of basins and tectonic belts exploring the new plate-tectonic concepts. Through the 1970s, a series of books was published containing the results. Many of these had their origins in the United States (e.g., Burk and Drake, 1974; Dickinson, 1974a; Dott and Shaver, 1974; Watkins et al., 1979) and Canada (Strangway, 1980; Miall, 1980). One of the more important of these contributions was a paper by Dickinson (1974b) which constituted the first comprehensive attempt to classify sedimentary basins of all types in terms of their plate-tectonic setting. This paper was particularly notable for the extensive treatment of arc-related basins, and was followed up by a more detailed paper on this subject (Dickinson and Seely, 1979) that remained the standard work on the subject for many years. This latter work was based in part on the recognition of a series of arc-related sedimentary basins within the Cordillera (Dickinson, 1976), especially the Great Valley basin of California, which has long served as a type example of a forearc basin (e.g., Ingersoll, 1978a, 1978b, 1979). Another important era in the field of basin analysis was initiated by the development of quantitative, geophysically based models of crustal subsidence, commencing in the late 1970s. Important early work was carried out by the Canadian geophysicist Walcott (1970a, 1970b, 1972), who explored the flexural properties of the crust through an examination of the effects of loading by the Laurentian ice cap and the Hawaiian volcanoes. Sleep (1971) studied the effects of the heating and elevation of continental crust at sea-floor spreading centers, and its subsequent cooling and subsidence during the formation of the classic Atlantic-type extensional margins (note, here, the significance of the Atlantic margin in becoming the type area for a particular class of continental margin). Keen (1979) confirmed Sleep’s ideas with reference to the continental margins of Nova Scotia and Labrador. However, there remained a serious space problem. Unless significant erosion of the updomed crust occurs, the crust will simply subside to its original level following the thermal event, and no space for sediments will have been created. Isostatic considerations show that only about half the thickness of crust removed by thinning or erosion can be replaced by sediment deposited up to sea level, a condition that cannot possibly explain the case of major Atlantic margin-type basins where sediment thicknesses of 5–10 km are typical. The main breakthrough in the development of a modern extensional-margin basin model was made by the Cambridge geophysicist McKenzie (1978), based in part on his studies of the subsidence of the Aegean Sea. This classic paper introduced the concept of crustal stretching and thinning during the heating event, and showed quantitatively how this could account for the subsidence history of Atlantic-type margins. Many of the important early tests of this model were carried out on the Atlantic margin of the United States. Stratigraphic data were obtained from ten Continental Offshore Stratigraphic Test (COST) wells drilled on the continental shelf off New England between 1976 and 1982, and led to the development of formal backstripping procedures (Watts and

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Ryan, 1976; Steckler and Watts, 1978; Watts, 1989) and to simple computer graphic models of subsiding margins (Watts, 1981) that were very useful in illustrating the development of the basic architecture of Atlantic-type margins. An important modification of the McKenzie model was to recognize the importance of simple shear during continental extension, as expressed by through-going extensional crustal detachment faults (Wernicke, 1985). This style of crustal extension was first recognized in the Basin and Range Province of Nevada, and was suggested by preliminary seismic data from the facing continental margins of Iberia and the Grand Banks of Newfoundland (Tankard and Welsink, 1987). The North Sea basin is the best studied rift basin, and has provided many insights regarding subsidence styles and structural geology (White and McKenzie, 1988, and many subsequent papers by White). Much of the work on the relationship between extensional tectonism and sedimentation was also initiated by British geologists, such as M. R. Leeder, and R. L. Gawthorpe (e.g., Leeder and Gawthorpe, 1987; Leeder et al., 1996). Turning to the other major class of sedimentary basins, those formed by flexural loading of the crust, it was Barrell (1917) who was the first to realize that ‘‘the thick nonmarine strata of the Gangetic plains accumulated in space made available by subsidence of the Indian crust beneath the mass of thrust plates of the Himalayan Range’’ (Jordan, 1995). Price (1973) revived the concept of regional isostatic subsidence beneath the supracrustal load of a fold-thrust belt that generates the marginal moat we now term a ‘‘foreland basin’’ (a term introduced by Dickinson, 1974b), based on his work in the Southern Canadian Cordillera. It is clear that the crust must have mechanical strength for a wide foredeep, such as the Alberta basin or the Himalayan foreland basin, to be created. The classic architecture of a foreland basin is defined by the isopachs of the sediment fill, which is that of an asymmetric lozenge, with a depocenter adjacent to the location of the crustal load, tapering along strike and also thinning gradually away from the orogen towards the craton. Two major developments contributed to our current understanding of these basins. Firstly, exploration drilling and reflection-seismic data led to an understanding of the structure and dynamics of the fold-thrust belts that border foreland basins and, during uplift, provide much of their sediment. Secondly, a growing knowledge of crustal properties permitted the development of quantitative models relating crustal loading, subsidence, and sedimentation. A significant development during the 1960s and 1970s was the elucidation of the structure of the fold-thrust belts that flank many orogenic uplifts and clearly served as the source for the clastic wedges referred to above. McConnell (1887) was one of the first to emphasize the importance of thrust faulting and crustal shortening in the formation of fold-thrust belts, based on his work in the Rocky Mountains of Alberta. As noted by Berg (1962), the mapping of faults in the Rocky Mountains of the United States and their interpretation in terms of overthrusting became routine in the 1930s. However, as his paper demonstrates, seismic and drilling data available in the early 1960s provided only very limited information about the deep structure of thrust belts. The release of seismic exploration data from the Southern Rocky Mountains of Canada by Shell Canada led to a landmark study by Bally et al. (1966) and set the stage for modern structural analyses of fold-thrust belts. A series of papers by Chevron geologist Clinton Dahlstrom, concluding with a major work in 1970 (Dahlstrom, 1970), laid out the major theoretical principles for the understanding of the thrust faulting mechanism. Jack Oliver was the first to suggest that reflection-seismic data could be used to explore the deep structure of the crust (Oliver et al., 1976). The concept was simple. Traditional industry seismic had been collecting the first six seconds of reflection energy; Oliver suggested extending the recording time to at least 15 seconds, which would continue the penetration to the base of the continental crust. His first experiment, in Texas (Oliver et al., 1976) identified the base of the Phanerozoic sedimentary section at 3.8 seconds two-way travel time, and the base of the crust, at the Moho, at about 12 seconds. He formed a group to establish the ‘‘Consortium for Continental Reflection Profiling,’’ (COCORP) headquartered at Cornell University. Use was made of the ‘‘Vibroseis’’ technique (sound vibrations) as the energy source, instead of dynamite. The deep seismic work of COCORP very quickly began solving longstanding problems of structural interpretation (Oliver, 1980, 1982). First up was a long-standing controversy regarding the deep structure of fault-bounded ‘‘block uplifts’’ or ‘‘basement uplifts,’’ including most of the Laramide structures of the US Rocky Mountains, the Boothia Uplift of Arctic Canada, the Wichita Uplift of Oklahoma, and many other comparable features. Until the 1940s it had been common practice to interpret these uplifts as being bounded by near-vertical basement faults, such a structural style being referred to informally as ‘‘thick-skinned,’’ in contrast to the ‘‘thin-skinned’’ style of thrust-fault belts. Berg’s (1962) synthesis of data from several of the major Laramide structures, including the Wind River Uplift, demonstrated that at least near the surface these were thrust faults, but not until the advent of deep-crustal seismic reflection surveying did it become clear that many of these faults extend at relatively high angles to the base of the crust (Smithson et al., 1978; Brewer et al., 1983). COCORP data also demonstrated the several hundred kilometers of overthrusting beneath the Appalachian orogen (Cook et al., 1979). At shallow depths, all of these structures are now known to have the character of thrust faults, and, as discussed later, thin-skinned tectonism has taken on a new meaning in certain orogenic belts. Beaumont (1981) and Jordan (1981) were the first to propose quantitative flexural models for foreland basins, constraining the models with detailed knowledge of the structure and stratigraphy of the studied basins. Their

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data were derived from Alberta, and the US Rocky-Mountains basin, respectively. Their papers became the standard references on the subject. Beaumont’s (1981) two-dimensional numerical model for the Alberta basin was later elaborated into a three-dimensional model and applied to the Appalachian foreland basin (Quinlan and Beaumont, 1984; Beaumont et al., 1988). The Appalachian orogen, and its extension around the southern margin of the continent as the Ouachita orogen, provided the basis for another important basin model, that of the remnant ocean basin. Tectonic reconstructions had indicated that the paleo-Atlantic (Iapetus) ocean had closed progressively from north to south, commencing with the Taconic (Ordovician) orogeny of the northern Appalachians, and ending with the Ouachitan orogeny (Mississippian-Pennsylvanian) of the southern continental margin, extending from Arkansas to northern Mexico. Distinctive patterns of sedimentation characterized the evolution of the remaining ocean basin, the deformed fill of which became the Ouachitan orogen (Miall, Chapter 8). Graham et al. (1975) showed how this pattern of events could be an important component of the ocean-closing phase of a Wilson Cycle. An important synthesis of basin studies was published by Bally and Snelson (1980). Most of the early work on extensional basins referred to here had appeared by this time, but not the foreland-basin models of Beaumont and Jordan. Basin terminology was still in a state of flux in the early 1980s. Bally and Snelson (1980) used such terms as episutural and perisutural basins, but these terms did not become widely used. Bally and Snelson (1980) also amplified an earlier concept proposed by Bally (1975), referring to ‘‘A-subduction,’’ the overriding of basement by a fold-thrust belt at the margin of a foreland basin. One purpose of introducing this term was to compare and contrast with ‘‘B-subduction’’ (also termed the Benioff zone), which is the term used to describe the descent of oceanic crust at the trench in front of an arc. A-subduction was thought to involve ‘‘substantial disposal at depth of possibly an earlier attenuated continental substratum’’ (Bally and Snelson, 1980, p. 20). Subsequent seismic work, to which we refer later, has since shown that this is not, in fact, a good description of what happens in such deformation zones, and the term A-subduction has, likewise, not survived. A few of the early COCORP lines were referred to by Bally and Snelson (1980) and had already helped to solve a long-standing controversy about the structural styles of basement uplifts, as we discussed above. COCORP and Lithoprobe data later gave a new meaning to thin-skinned tectonism (Bally’s A-subduction), which we discuss later. Basin classifications of Halbouty et al. (1970), Dickinson (1971), Klemme (1975), and Bally and Snelson (1980) reflect an incomplete absorption of plate tectonics into the analyses being carried out by basin analysts; but by the mid-1980s, basin classifications in terms of plate tectonics were beginning to stabilize. New classifications were offered by Kingston et al. (1983), Miall (1984), Mitchell and Reading (1986), Klein (1987), and Ingersoll (1988). All these studies cited the classification of Bally and Snelson (1980) and, while they all differed substantially from that work, the differences between them were not that large. These basin concepts and models formed the basis for much of the analysis of North American geology undertaken during the DNAG period of the 1980s and early 1990s. A later article by Bally (1989), in his introductory article in volume A, reveals how far basin analysis had advanced in the intervening decade. He compared and contrasted some of the major North American basins at length in order to extract and illustrate the major themes of these two basin types. In his conclusions (p. 439) he pointed out the need for more threedimensional seismic data in order to document the details of basins and their relationship to basement. This was, of course, exactly the kind of data that was in the process of being provided by COCORP and Lithoprobe. At about this time, Allen and Allen (1990) published their comprehensive textbook ‘‘Basin analysis: Principles and applications’’ which focused on the physics of crustal evolution and the quantitative approaches that had evolved to explore the relationships between sedimentation and tectonics. This landmark publication set a new standard for teaching and research in basin analysis over the following decade. Heller et al. (1988) and Heller and Paola (1992) examined the common assumption that the coarse, commonly conglomeratic top to stratigraphic pulses in foreland basins record the time of maximum tectonism. They termed this the ‘‘syntectonic model’’ of sedimentation. The term ‘‘clastic wedge’’ had been suggested by King (1977) for lobes of sediment extending away from orogenic sediment sources, and he provided a classic illustration of the interpreted syntectonic turbidite and fluvial clastic wedges of the Appalachian orogen. But the ‘‘syntectonic’’ interpretation may be incorrect. It depends (among other factors) on which response to thrusting and supracrustal loading is most rapid, crustal flexure and subsidence or the development of an organized drainage net to carry detritus from uplifted source areas. In another scenario, which Heller and his colleagues termed the ‘‘antitectonic model,’’ the time of thrusting and crustal loading may be one of basin deepening, with the deposition of fine-grained deposits, the coarsening-upward succession recording the gradual establishment of drainage nets, and the building out of coastal-plain depositional systems, possibly long after the triggering pulse of tectonism has ceased. Other scenarios are possible, depending on the flexural response to loading (which depends on crust thickness and age), rates of erosion and transportation (climate and source area variables), and local paleogeography, which governs the ultimate distribution of the detritus (e.g., removal from the basin by through-going rivers or entrapment by basement topography). The relationships among these variables are

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complex, and researchers should be aware of drawing simplistic conclusions (Jordan et al., 1988; Heller and Paola, 1992). Work both in Europe and North America addressed the relationship between the dynamics of fold-thrust and nappe evolution, crustal loading and sedimentation. Diffusion concepts were introduced to simulate the physics of sediment transport. Sinclair et al. (1991) and Sinclair and Allen (1992) developed models for the evolution of the Swiss molasse basin. Riba (1976) and Anado´n et al. (1986) described the extraordinarily well-preserved syndepositional intraformational unconformities developed at the margin of the Ebro Basin, Spain. Meanwhile, P. G. DeCelles and colleagues were exploring the relationships between fold-thrust-belt development and syntectonic sedimentation in the Cordilleran Mountains of the western United States and the Himalayas (e.g., DeCelles et al., 1991, 1998). A very influential numerical two-dimensional model for foreland basins was developed by Flemings and Jordan (1989) which explored the relationships between subsidence, sedimentation, and eustatic sea-level change. This approach was later applied more generally to other basin settings, such as extensional continental margins ( Jordan and Flemings, 1991). Another useful contribution was that of DeCelles and Giles (1996), who showed that foreland basins should be subdivided into four parts, each showing a distinctive history. Since the work of Walcott (1970a, 1970b) it had been known that a flexural load on the crust is accompanied by an uplift, or forebulge, several hundred kilometers out from the point or locus of load, and it had also been recognized that there are commonly minor sedimentary basins within the fold-thrust belt, where sediment is trapped between emergent structural culminations. DeCelles and Giles (1996) formalized the various components of foreland-basin systems into the so-called wedge-top, foredeep, forebulge, and backbulge basins, and discussed the patterns and mechanisms of subsidence within these partitions. Uplift and subsidence of the forebulge typically take place in the direction opposite vertically to that of the foredeep, in response to crustal loading and relaxation, respectively. A detailed study of part of the Alberta basin by Catuneanu et al. (1997b) demonstrated that this process occurs over a frequency in the 104–107-year range. DeCelles and Giles (1996) and Catuneanu et al. (1997a) also discussed a concept developed earlier by Cross (1986) and Mitrovica et al. (1989), that subsidence of foreland basins is influenced by the thermal history of the underlying mantle. A cold, downwelling mantle current above the downgoing oceanic plate at depth results in what Catuneanu et al. (1997a, 1997b) called the ‘‘dynamic load,’’ which may increase the accommodation in a foreland basin by several hundred meters and its width by hundreds of kilometers. The magnitude of the dynamic load during much of the Cretaceous is the reason why the forebulge is so difficult to map over much of the Alberta basin (Miall et al., Chapter 9). Ettensohn (Chapter 4) provides a useful discussion of these concepts as applied to the Appalachian basin, including a modern version of what used to be called the ‘‘geosynclinal cycle’’ (Figure 6 in Chapter 4). Beginning in the late 1980s, Cloetingh et al. (1985) and Karner (1986) added a very important component to basin tectonics, the elucidation of the importance of intraplate or in-plane stress transmitted from extensional or contractional plate margins. Significant vertical motions of the crust may be explained by this mechanism, which also explains the occurrence of significant intraplate earthquakes. The stratigraphic implications of intraplate stress are significant (Cloetingh, 1988). Zoback (1992) reported on a project to document these stress patterns on a global basis. Dynamic subsidence of foreland basins is now known to be part of a broader process which goes by the term epeirogeny. Stratigraphers specializing in the study of continental interiors (e.g., Sloss and Speed, 1974) have for a long time appealed to this process, first described and named by Gilbert (1890). Epeirogeny is ‘‘a form of diastrophism that has produced the larger features of the continents and oceans, for example plateaux and basins, in contrast to the more localized process of orogeny, which has produced mountain chains.’’ (Bates and Jackson, 1987) The definition goes on to emphasize vertical motions of the earth’s crust. Modern studies of the thermal evolution of the mantle, supported by numerical modeling experiments, have provided a mechanism that explains the long-term uplift, subsidence, and tilting of continental areas, especially large cratonic interiors beyond the reach of the flexural effects of plate-margin extension or loading (Mitrovica et al., 1989; Gurnis, 1988, 1990, 1992; Burgess and Gurnis, 1995; Burgess et al., 1997; Burgess, Chapter 2). These studies have shown that the earth’s surface is maintained in the condition known as dynamic topography, reflecting the expansion and contraction of the crust resulting from thermal changes in the underlying mantle. Dynamic uplift occurs when mantle heat is trapped beneath the crust of a large supercontinent, and the resulting uplift can help to explain the occurrence of low sea levels during times of supercontinent assembly (e.g., during the Late Paleozoic–Triassic Pangea period). Is there a reason why the early developments in quantitative basin models, in plate-tectonic basin classifications, and in more recent research in basin dynamics have been largely North American in origin? Perhaps the basically simple architecture of the North American plate has something to do with this. The continental plate is bounded on not one, but two sides, by a simple arrangement of craton-foredeep-fold/thrust belt-orogenic collage-classic margin-ocean. Comparable two-dimensional cross-sections through these margins can be drawn in many places (e.g., see Figure 4 in Chapter 1). The quantitative models of continental collision and foreland-basin development proposed by Stockmal et al. (1986) developed the theme of the overriding of an

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earlier extensional margin by a colliding orogen, superimposing a foreland basin over the older, passive continental-margin wedge. This model applies to much of the Appalachian and Ouachitan margin (Lavoie, Chapter 3; Ettensohn, Chapter 4; Miall, Chapter 8), to the Antler orogen of California-Nevada (Ingersoll, Chapter 11), and to the early stages of the Cordilleran orogen of central British Columbia (Ricketts, Chapter 10). Much of European and Asian geology is far more complex, and perhaps it is not surprising that some of the major developments made there reflect this complexity. For example, the leading British sedimentary geologist, Harold Reading, focused attention on the importance of crustal strike-slip faults and the subtleties of lateral displacement within orogens (Reading, 1980). Much of the interesting work on orogenic suture zones from this period is also European and Asian in origin, such as the several papers on the Alpine-Himalayan suture in the book edited by Coward and Ries (1986), the review article by Dewey (1977), and the studies of the Tethyan suture by Sengo¨r (1976, 1986, and many other articles). Other significant exceptions to my gross generalization include the work on foreland basins by Sinclair and co-workers, referred to above, and the important set of papers on foreland basins compiled and introduced by Allen and Homewood (1986), and the introduction of the concept of intraplate stress by S. Cloetingh (references provided earlier). There are few regions in North America that compare on the regional scale with the complexity of the Hercynides in western Europe (I still recall the then graduate student, John Bradshaw, demonstrating multiple generations of fold development in Paleozoic strata of Brittany, during an undergraduate field trip in 1964), although complexities abound within the accreted terranes of the Appalachian and Cordilleran orogens, the unraveling of which required the introduction and elaboration of the terrane concept, another largely North American development, to which we turn later. For an example of Phanerozoic complexity one might turn to the region of the Ancestral Rockies in the southwest part of the continent (Blakey, Chapter 7). Paleogeographic and tectonic reconstructions based on early plate-tectonic concepts, including the Wilson Cycle, assumed that earth’s history had been dominated by the movement of large continental blocks, but during the 1970s this began to be seen as overly simplistic. Schermer et al. (1984, p. 107) wrote: ‘‘A growing body of geological and geophysical evidence demonstrates that large translational and rotational displacements have occurred within and between tectonic provinces of the eugeosynclinal belt. A natural outgrowth of plate-tectonic theories — the concept of terrane analysis — recognizes that the diverse fragments of vastly different geologic histories cannot be assumed to have genetic relationships in time and space, and therefore are ‘suspect.’’’ Schermer et al. (1984, p. 110) defined a terrane as ‘‘a fault-bounded geologic entity of regional extent that is characterized by a geologic history different from that of neighbouring terranes.’’ Mapping of the North American Cordillera, particularly in Alaska and California, led to the recognition of terranes for which paleogeographic affinities and tectonic history were very difficult to unravel. In many cases all that could be said for certain was that the geology did not appear to be related in any simple way to that of adjacent terranes. From this came the terms ‘‘suspect terrane’’ and ‘‘allochthonous terrane,’’ both terms indicating movement and emplacement from sources unknown. Some early plate-tectonic syntheses, such as the reconstruction of the Alpine orogen by Dewey et al. (1973), had indicated that suture belts might be composed of ‘‘small continental fragments,’’ or ‘‘microplates,’’ but these reconstructions attempted to trace a coherent tectonic history for each continental fragment. Dewey et al. (1973, p. 3,140) suggested that ‘‘Possibly there were, in the past, numerous other plates that have been welded together in an all but indecipherable maze;’’ but they thought not. They considered that they had been able to identify the origins and track the kinematic history of all the small components of the Alpine orogen, albeit with some difficulty and ambiguity. Dickinson (1976) attempted a similar type of analysis for the North American Cordillera, based on the evolution of a simple Andean-type arc system, but several indicators of problems with this type of analysis were already apparent. The geology of parts of the Cordillera is particularly complex, and this was the source of a perhaps uniquely North American contribution to the plate-tectonic interpretation of complex orogens. The ‘‘terrane’’ concept is attributed to Irwin (1960, 1972), based on his work in the northern Coast Range and Klamath Mountains of California, where the geology posed many challenges. One of the earliest indications that particular continental fragments (‘‘terranes’’) might be far traveled came from the work of Monger and Ross (1971), who identified exotic faunas in the Cache Creek terrane of central British Columbia that showed Asian affinities. Wrangellia could probably be designated as the type example of a terrane. This terrane consists of five separate blocks, originally thought to have derived from a Late Paleozoic oceanic plateau covered by a limestone platform somewhere in the southern Panthalassa Ocean, and now dispersed in five separate pieces between Oregon and the Wrangell Mountains of southern Alaska ( Jones et al., 1977). By the end of the 1970s, terrane analysis of the Cordillera and other complex orogens had become the new orthodoxy (Coney et al., 1980; Ben-Avraham et al., 1981). However, the concept of a distinct philosophy and methodology of ‘‘terrane analysis’’ is by no means universally popular. Practitioners of the new methodology (e.g., Schermer et al., 1984) claimed that the method is objective because terrane definition does not depend on the identification of a particular platetectonic scenario for each terrane. However, this can lead to the definition of a new terrane wherever geological relationships are obscure. By retreating to a somewhat descriptive approach, Sengo¨r and Dewey (1991)

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complained that the method ‘‘takes the responsibility of interpretation of earth history off the shoulders of geologists’’ by removing the obligation to develop and incorporate logical, testable models for the plate-tectonic evolution of an orogen. To some extent this argument boils down to a controversy between empiricism and model building and the approaches preferred by practitioners, a controversy comparable to that described elsewhere with respect to the history of stratigraphy (Miall, 2004; Miall and Miall, 2004), and discussed later. Chapters in this book that deal with terrane geology (see, in particular, Chapter 10, by Ricketts), clearly show how the methods of terrane analysis have adapted some very standard geological techniques for the purpose of exploring the docking and tectonic evolution of terranes and terrane-related sedimentary basins. For example, minimum dates for terrane amalgamation may be derived from the ages of cross-cutting or ‘‘stitching’’ plutons and by the age of sediments constituting ‘‘overlap assemblages,’’ that is, sedimentary units containing detritus derived from one terrane and deposited on the adjacent terrane. Apparent polar-wander data derived from paleomagnetism have been fundamental to terrane analysis, although these data are commonly associated with major controversy (e.g., the ‘‘Baja-BC’’ controversy, referred to later). There would no longer seem to be the need for debate about the competing theories of traditional plate-tectonic analysis versus terrane analysis, but there is plenty of work to do to sort out remaining problems of terrane history in many parts of the world. By the late 1980s and early 1990s, towards the conclusion of DNAG, and aided by major projects, such as COCORP and Lithoprobe, earth scientists developed a new confidence about the geology of the North American continent. For example, Allmendinger et al. (1987) suggested that deep seismic reflection data from the continental crust indicated four basic reflection fabrics, one corresponding to the craton, with a thin sedimentary cover, a second zone developed in areas of thin-skinned tectonism beneath fold-thrust belts, a third zone defined for the area of ramp-like reflections of steeply-dipping detachment faults, and the fourth zone corresponding to the areas of strongly layered lower crust that develop in areas of late-orogenic crustal extension. Williams et al. (1991) suggested that the time was right for the construction of new thematic maps drawing together a range of continent-wide facts relating to the tectonic evolution of the continent. They highlighted such details as the major tectonic elements, including terranes, times of collision, suturing, and accretion. Percival et al. (1999) revealed an even more sophisticated level of understanding about crustal processes. They proposed four models of collisional orogen that reflected such differences as the degree of homogeneity and resulting strength of the overriding plate, citing several global examples. They also referred to the increasing knowledge obtained from the coupling of geochemical with geophysical studies of the deep crust and underlying mantle, emphasizing the new knowledge being gained about the recycling of crustal materials at subduction and suture zones. A new appreciation was expressed of the importance of decoupling between crust and mantle. The extensive crustal shortening involved in the successive arrival and obduction of a series of terranes onto the western continental margin of North America led to processes called tectonic wedging and delamination (Price, 1986). Lithoprobe data show that the entire upper crust across the Cordillera of British Columbia has become separated — delaminated — from lithospheric basement (Cook et al., 1995) and translated hundreds of kilometers cratonward across virtually flat de´collement zones. A similar process also affected the Appalachian orogen (Sacks and Secor, 1992; Gibling et al., Chapter 6), and is likely a structural characteristic of all complex collisional orogens. Beneath Newfoundland, Lithoprobe data reveal the underthrusting of the two major accreted terranes, Ganderia and Avalonia, beneath the Laurentian craton, resulting from aborted subduction (van Staal, 2005; Percival et al., 2004). As Gibling et al. (Chapter 6) noted, a full comprehension of complex orogens such as the Cordilleran and Appalachian requires and, indeed, can now be attempted, using a ‘‘lithosphere-scale’’ approach to tectonic analysis and basin studies, employing combinations of all modern geological and geophysical tools. Percival et al. (2004) returned to the theme of orogen classification, proposing that a distinction could be made between orogens developed by simple subduction (e.g., the Andes), orogens involving terrane accretion (e.g., the North American Cordillera), and those involving major plate collisions (e.g., the Himalaya). Clearly, individual orogens might evolve from one of these stages to another. Also considered were differences between orogens relating to the state of the pre-existing continental margins, the time gap between rifting and collision, which determines, among other factors, the thickness of rift sediments and the style of magmatism.

1.2. Stratigraphy and sedimentology This is not the place for a dissertation on the history of stratigraphy and sedimentology in North America. However, some points are emphasized here, based on more lengthy treatments of the subject by Miall (2004) and Miall and Miall (2004). It can, once again, be argued that several of the key developments are North American in origin, and reasons can be offered for this. In what became a classic paper Barrell (1917) explored the origins of stratigraphic units in terms of depositional processes. Aided by the new knowledge of the earth’s radiogenic heat engine and a growing understanding of sedimentary processes, Barrell worked through detailed arguments about the rates of

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sedimentation and the rates of tectonism and of climate change. He constructed a diagram showing the ‘‘Sedimentary Record made by Harmonic Oscillation in Baselevel’’ which is remarkably similar to diagrams that have appeared in some of the Exxon sequence model publications since the 1980s that show the relationship between subsidence, base level change, and the creation and removal of sedimentary accommodation. Barrell’s classic paper begins in this way: ‘‘Nature vibrates with rhythms, climatic and diastrophic, those finding expression ranging in period from the rapid oscillation of surface waters, recorded in ripple mark, to those long-defended stirrings of the deep titans which have divided earth history into periods and eras.’’ Barrell was aware of climatic cycles and discussed what we would now call orbital-forcing mechanisms (e.g., the ‘‘precession cycle of 21,000 years’’). He discussed the major North American orogenic episodes and their influence on the broad patterns of stratigraphy, referring (p. 775) to the rise and ebb of sea level, which ‘‘pulsates with the close of eras, falling and then slowly rising again,’’ as ‘‘the most far-reaching rhythm of geologic time.’’ However, the main focus of this important paper is on attempts to establish the rates of geological processes and the measurement of the length of geologic time, given the new impetus to the study of this problem provided by the discovery of radioactivity. He refers to ‘‘diastrophic oscillation,’’ but only from the understanding such a process may provide for the interpretation of the stratigraphic record, not as a fundamental mechanism to be used as a basis for the definition of geologic time. Since the late 1970s, stratigraphic practice has been dominated by the methods of sequence stratigraphy. Sequences are defined as ‘‘unconformity-bounded successions of strata’’ (Sloss, 1963; Vail et al., 1977). Their value as stratigraphic concepts is threefold: they have an internal architecture that is commonly quite predictable; they have wide regional extent, which can be reconstructed to yield wide-ranging stratigraphic and paleogeographic syntheses; and their architecture and/or composition yield clues as to large-scale basin controls, including tectonism, sea-level change, and climate change. The subdivision of the stratigraphic record into sequences was primarily an American idea. Blackwelder (1909) published a diagram showing the extent of some major unconformities and their intervening stratal packages across the North American interior that in many ways anticipated the classic diagram of Sloss (1963). During the 1930s and 1940s, the concept of widespread regional packages of strata became commonplace in North America, as reflected by the work of the distinguished American petroleum geologist A. I. Levorsen, who used such terms as ‘‘natural groupings of strata on the North American craton’’ and ‘‘successive layers of geology in the earth, each separated by an unconformity’’ (Levorsen, 1943, p. 907). Through the first half of the twentieth century and, indeed, until the post-WWII phase, when Sloss was beginning his work, much of the interior and western regions of North America were still in what petroleum geologists would call the ‘‘frontier’’ stage of exploration, for which regional syntheses would be of considerable utility. Europeans, by contrast, had been documenting the regional stratigraphy of their continent for more than one hundred years. Perfection of the minutiae of correlation was a central focus of stratigraphic research, and attempts to define broad cyclic patterns in the European stratigraphic record (e.g., Ramsbottom, 1973, 1977), were strongly resisted in some quarters (e.g., George, 1978). The focus of American work was, therefore, different and it did not help that the ‘‘standard’’ chronostratigraphic divisions of geologic time, that had so painstakingly been worked out mainly by Europeans (Berry, 1987) did not seem to work in North America. Many of the ‘‘natural breaks’’ in the stratigraphic record appeared to be in the ‘‘wrong’’ places. We now know, of course, that this is because natural breaks (regional unconformities) are largely the product of tectonism, which is at most continental in scope, rather than intercontinental. Sequence stratigraphy encourages a large-scale approach to paleoenvironmental analysis. This is encapsulated by the terms ‘‘depositional system’’ and ‘‘systems tract.’’ The concept of depositional systems and the idea of largescale, repeated cycles of sedimentation, termed ‘‘depositional episodes,’’ both derive from work along the Gulf Coast of the United States, where depositional systems are developed on a large scale. The seminal work of Frazier (1974) is of particular importance in the development of sequence stratigraphy, although this is not properly credited in most review literature. The depositional-episode concept was developed by Frazier (1974) to describe and explain the cycles of post-glacial sedimentation on the Mississippi delta (see Galloway, Chapter 15), and served as a major inspiration for the analysis of sequences in terms of depositional systems tracts by Vail et al. (1977), Posamentier and Vail (1988), Posamentier et al. (1988), and Van Wagoner et al. (1990). The ‘‘facies-cycle wedge’’ concept of White (1980) related large-scale facies relationships and oil and gas plays to transgression and regression. The ideas were popular for awhile, but were rapidly supplanted by sequence stratigraphy, especially when sequence stratigraphic concepts were applied to outcrop and well data by the Posamentier, Vail, and Van Wagoner articles cited above. A quite separate branch of cyclic sedimentology had grown up around the Late Paleozoic mid-continent ‘‘cyclothems,’’ described briefly elsewhere in this book (Miall, Chapter 8). The first description of these is attributed to J. A. Udden, who published his findings in 1912 (see Langenheim and Nelson, 1992). In 1926, the Illinois Geological Survey began a stratigraphic mapping study of these deposits, under the direction of J. Marvin Weller. Mapping by Weller and his colleagues was the first to demonstrate that these cyclothems are present over

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much of the continental interior of the United States (Weller, 1930; Wanless and Weller, 1932). Weller (1930) suggested diastrophism as the cause of the cycles, but a different mechanism was proposed a few years later. ‘‘It happens that there is abundant evidence of the existence of huge glaciers in the southern hemisphere during the very times when these curious alternations of deposits were being formed. A relation between these continental glaciers and the sedimentary cycles has been proposed recently by the writers’’ (Shepard and Wanless, 1935). The cyclothem concept has been enormously influential in the study of cyclic sedimentation since that time, but cyclothems represent an unusual set of circumstances, control by high-frequency glacioeustatic sea-level changes, a mechanism which cannot be invoked for many other types of cyclic succession. Van Siclen (1958) developed a sedimentological model for the development of clastic and carbonate platform margins formed under the influence of high-frequency sea-level change — his rocks were the cyclothems at the southern platform margin in Texas (see Miall, Chapter 8) — a model which anticipated modern sequence stratigraphy by some thirty years, but this work is not referenced in any of the key modern theory or review papers. The plate-tectonic revolution has encouraged some earth scientists to ‘‘think big’’ about sedimentation and tectonics. For example, Dickinson (1988) speculated about the control of plate movements and orogeny on uplift and erosion, the development of continental-scale drainage systems, and the patterns of sediment dispersal. He offered a single speculative example, the possible fate of detritus eroded from the Appalachian orogen. We come back to his example in the next section. The debate about the importance of eustatic sea-level change as a control on regional stratigraphy is discussed elsewhere (Miall and Miall, 2001).

2. North American Geology There are unusual, even unique features about the sedimentary geology of the North American continent. The Mesozoic Western Interior Seaway (Miall, et al., Chapter 9) is one of the largest foreland basins on earth, extending north–south over about 351 of latitude, from Texas to the Northwest Territories (Figure 1 in Chapter 9), a distance of more than 3,000 km. The continental embankment represented by the Gulf Coast basin (Galloway, Chapter 15) is also one of the largest features of its kind on earth, extending from Oklahoma to the Sigsbee salt escarpment, a distance of close to 1,000 km, with sediment thicknesses in the range of 16 km over wide areas. Galloway (Chapter 15) notes that the gravity-driven tectonic structures of the Gulf Coast, driven by salt diapirisim, are amongst the most extensive and best described in the world. Pleistocene strata in the Gulf of Mexico are locally more than 7.6 km thick, as a result of some of the highest subsidence rates on earth, reflecting the huge influx of sediment from the Mississippi system (Worrall and Snelson, 1989; Ingersoll and Busby, 1995). The Atlantic margin of the continent, a continuous tectonic province that has undergone a coherent platetectonic evolution since the Late Triassic, is also enormous, extending a distance of 7,000 km from northern Baffin Island to southern Florida (Miall et al., Chapter 14). The ‘‘broken foreland basin’’ represented by the Laramide basins of the Rocky Mountain states (Lawson, Chapter 12), if not the only group of basins of this type in the world, is certainly the best known (Ingersoll and Busby, 1995). Basins formed in strike-slip or transform settings are also particularly well-known from the San Andreas system of western North America, but are not described in this book (see Nilsen and Sylvester, 1995).

2.1. Evolution of the craton The North American craton is not the only large craton amongst the earth’s continental plates, but is one of the largest, and almost certainly the best known, in terms of outcrop and subsurface data and the amount of study by earth scientists. Several major themes have emerged from the long-term study of this craton. By Late Precambrian time, the north American craton was almost entirely surrounded by extensional continental margins. Phanerozoic North America would have been like modern Africa and Australia in terms of the continental tectonic setting (Hoffman, 1989; see Chapter 1, Figure 3). The Precambrian–Paleozoic unconformity is a profound angular unconformity across much of the cratonic interior of North America (it is a conformable contact in some of the marginal basins, e.g., in the Appalachian basin; see Ettensohn, Chapter 4). From Newfoundland to Florida, the Precambrian–Paleozoic unconformity records the creation and erosional removal of an entire mountain range. It seems likely that the global scale of Grenville tectonism generated an orogen of the magnitude of the modern Himalayan ranges (Hoffman, 1991). The metamorphic grade of the Grenville rocks now exposed at the surface suggest that 25–30 km of the orogenic superstructure had been removed by early in the Phanerozoic, and the Cambrian–Ordovician transgression took place across a virtually flat surface (Anovitz and Essene, 1990; Hoffman and Grotzinger, 1993).

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The high sea levels that prevailed during the Ordovician–Devonian and again during the Cretaceous developed interior cratonic shelf seas that have no modern counterpart. Some depositional systems, such as some of the Lower Paleozoic carbonate units, and the Devonian evaporites of the Prairies, can be traced for hundreds, even thousands, of kilometers, and are difficult to interpret on the basis of modern analogues. The successive development of major collisional orogenic belts, first on the eastern and southern margins of the continent (tectonic phases one and two of Miall and Blakey, Chapter 1) and later on the western margin (phase three) had long-term effects on intraplate stress and the dynamic topography of the craton. The tectonic and sedimentary results are described in some detail by Burgess (Chapter 2). Earlier workers, such as Sloss (1988) and Bally (1989), had remarked on some of the more significant events. For example, Bally (1989, p. 425) noted that towards the end of the Silurian and the beginning of the Devonian period there was a change along the craton margins from an overall cratonic-miogeoclinal regime (Sauk and Tippecanoe sequences) to a dominantly foredeep regime (Kaskasia and younger sequences). Movements of Boothia Uplift in the Arctic and the Peace River and Transcontinental arches probably reflect intraplate contractional stresses imposed by Appalachian– Caledonian collisions (see also Miall and Blakey, Chapter 1; Burgess, Chapter 2; Miall, Chapter 5). During the Late Carboniferous there was a ‘‘profound reorganization of the craton’’ (Bally, 1989, p. 428) indicated by changes in isopach patterns in the southwest part of the craton (see Blakey, Chapter 7). This was a result of the Ouachita collision, which was essentially complete by the Mid-Pennsylvanian (Miall, Chapter 8). Another episode of major change was the Absaroka/Zuni unconformity (Early Jurassic) (Sloss, 1988; Blakey, Chapter 7), reflecting the beginning of phase three tectonism on the western continental margin.

2.2. Sediment fluxes Orogenic uplift and the tilting and warping of the craton have led to extended periods of subaerial erosion and some major fluxes of sedimentary detritus. As noted above, metamorphic grades of the Grenville indicate the removal of some 25–30 km of rock from this province prior to the Cambrian–Ordovician (Sauk) transgression. Where did this detritus go? Rainbird (1992), Rainbird et al. (1992, 1997), and Hoffman and Grotzinger (1993) suggested that much of it may have ended up contributing to the thick Neoproterozoic sedimentary wedges on the western continental margin. Detailed study of detrital zircons from sedimentary rocks of this age in the western Canadian Arctic indicated that 50% of them are of Grenville age. Rainbird and his colleagues proposed that a major west-flowing river system was established during the Late Proterozoic which transported this detritus some 3,000 km across the continental interior. Sedimentary evidence from the Canadian Arctic supports the existence of very large, deep rivers at this time (Rainbird, 1992). This was only the first of several episodes of large-scale sediment transport across the Laurentian interior and around its margin. Major drainages, on the scale of the modern Mississippi–Missouri–Ohio river system, and turbidite systems on the scale of the modern Bengal and Mississippi fans have developed at several times during the Phanerozoic. Major river systems evolved during times of continental exposure, corresponding to the unconformities between the Sloss sequences, but the direction of sediment transport and the location of sediment accumulation have changed significantly through the Phanerozoic (Miall, 2006). Neodymium geochemistry has been used to identify sediment sources for the Devonian clastic wedge of the Franklinian basin in the Canadian Arctic. The most likely source is the Caledonian orogen of eastern Greenland. Patchett et al. (1999) argued that uplift, erosion, and westward transport of this detritus commenced following Caledonian orogeny in the Early Silurian, with sediment undergoing recycling and further transport throughout the Silurian and Devonian, ultimately constituting one of the major sources of detritus for clastic units along the western continental margin as far south as Alberta. Ettensohn (Chapter 4, see Figure 27), following Archer and Greb (1995), suggested the presence of an ‘‘Amazon-scale drainage system that headed in the Canadian Shield’’ during the Early Pennsylvanian, carrying detritus southwestward through the Appalachian foreland basin. Patchett et al. (1999) speculated that the Ouachita turbidites, transported along the axis of the remnant ocean basin as the Appalachian collision progressed (Graham et al., 1975), were derived from the southern extension of this same collisional orogen. Much of this detritus ended up in the Marathon basin of west Texas (Gehrels et al., 2007). On the basis of coal moisture measurements and other measures of organic metamorphism, it is known that massive unroofing of the Appalachian and Ouachitan orogens occurred during the Late Paleozoic-Early Mesozoic. Removal of 4 km of sediment occurred within the undeformed Appalachian foreland basin; there was 7–13 km of cumulative erosion in the central Appalachians, and more than 13 km in the Ouachitas (Beaumont et al., 1988). Bally (1989, p. 430) said: ‘‘These amounts of erosion are stunning, and raise the question of the fate of the eroded material and how much of the unloading was due to as yet unrecognised tectonic unroofing.’’ Dickinson (1988) had suggested that much of the thick accumulations of Late Paleozoic and Mesozoic fluvial and eolian strata in the southwestern United States had been derived from Appalachian sources, and this was

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confirmed by the zircon studies of Dickinson and Gehrels (2003). Neither the location nor scale of specific river systems could be identified by this research, but given the distance of transport and the volume of sediment displaying these provenance characteristics that is now located on the western continental margin, it seems highly probable that large-scale river systems were involved. A westerly tilt of the craton during the Late Paleozoic and Early Mesozoic could at least in part be attributed to continental-scale thermal doming preceding the rifting of Pangea, followed by the more localized heating and uplift of the rift shoulder of the newly formed Atlantic rift system. Ettensohn (Chapter 4) suggested that some of the detritus was also shed eastward, resulting in the accumulation of more than 5 km of sediment in southern New England in Early Jurassic time. The major uplift and erosional unroofing took place during the Late Permian to Early Triassic, and by Late Triassic time ‘‘parts of the old orogen in New England had been exhumed to nearly the present erosional level’’ (Ettensohn, Chapter 4; citing isotopic studies of Dallmeyer, 1989). During the Cenozoic, available evidence suggests that big rivers in the North American continental interior were flowing the other way, towards the east, with the deposition of thick clastic wedges along the continental margin off the east coast, including the continental shelf and slope of Baffin Island and Labrador (McMillan, 1973; Duk-Rodkin and Hughes, 1994). It is known that at least 2 km of sediment has been stripped from the Cretaceous–Cenozoic Cordilleran clastic wedge in Alberta (Beaumont, 1981) and remnants of the last clastic pulse derived from uplift of the Cordillera are preserved as patches of Eocene–Oligocene gravels across southeastern Alberta and southern Saskatchewan (Leckie, 2006). The reversal in regional drainage is attributed to the formation and uplift of the Cordilleran orogen, with a likely contribution, also, from dynamic topography effects related to mantle thermal systems that were re-ordered as Pangea broke up and the North American continent began to be carried westward. As shown by Fisher and McGowen (1967) the Mississippi–Missouri–Ohio system and other rivers draining into the Gulf Coast were established early in the Cenozoic following the Laramide dismemberment and uplift of the Western Interior Seaway (Lawton, Chapter 12). The enormous clastic wedges in the Gulf of Mexico basin indicate the existence of the major river draining out into the Gulf since at least the Eocene, with major clastic episodes reflecting pulses of uplift and erosion of the Cordillera (Galloway, Chapter 15). The integrated postMid-Oligocene river system draining the Cordillera was then interrupted in the north by the Late Cenozoic glaciation, which blocked eastward drainage across the Canadian Shield and ultimately led to the development of overflow channels that evolved into the modern Mackenzie River, draining into the Arctic Ocean (Duk-Rodkin and Hughes, 1994). Rittenour et al. (2007) provided a detailed assessment of the evolution of the Mississippi river valley in response to glaciation and deglaciation.

2.3. Other effects of intraplate stress Intraplate stresses can now be seen to be the cause of many local tectonic events that cannot readily be assigned to the regional plate-tectonic chronology. For example, uplift of the Peace River Arch (Alberta) and Ancestral Uinta Uplift (Utah) during the Mid-Devonian (Miall, Chapter 5), and several episodes of movement of the Transcontinental Arch (Burgess, Chapter 2) may be explained as the product of the transmission of contractional stresses from craton-margin orogenic episodes across the continental interior. Boothia Uplift and several other small tectonic elements in Arctic Canada had long been attributed to enigmatic block uplifts of unknown cause (e.g., Kerr, 1977), but can now be understood as a product of intraplate stresses transmitted from the Caledonian collision between Greenland and Scandinavia (Miall, 1986; Okulitch et al., 1986). Galloway (Chapter 15) documents ‘‘Laramide’’ thrusting in Gulf of Mexico basin, which is well outside the area affected by Laramide tectonism (Lawton, Chapter 12). Also in this chapter, Galloway notes that most unconformities in the Gulf Coast clastic wedges, notably the prominent Mid-Cretaceous unconformity, are not a result of sea-level fall but reflect changes in the crustal stress regime across the continent. Cenozoic clastic pulses there reflect Cordilleran and Appalachian tectonism. Poag and Sevon (1989) compiled isopach maps of the post-rift sedimentary record along the US Atlantic margin, and interpreted these in terms of the shifting pattern of denudation in and sediment transport from the Appalachian and Adirondack Mountains. It seems likely that many of the major existing rivers that presently carry sediment into this region have been in existence since the Mesozoic. Poag and Sevon (1989) were able to show many changes with time in the relative importance of these various sediment sources, and it is suggestive that several of these changes occurred at times of change in Atlantic plate kinematics. The sea-floor spreading record reveals at least fourteen changes in plate configuration in the central Atlantic since the Mid-Jurassic (Klitgord and Schouten, 1986). Seven of these changes, at 2.5, 10, 17, 50, 59, 67, and 150 Ma, correspond to times when the major sediment dispersal routes from the Appalachian Mountains to the continental shelf underwent a major shift (data from Poag and Sevon, 1989), suggesting that the pattern of uplift (and subsequent erosion) of the coastal mountains is directly affected by changes in intraplate stresses transmitted from the adjacent Atlantic oceanic plate.

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3. Remaining Problems A review of recent work on the geology of the North American continent suggests that most of the larger problems have been resolved. There is no longer serious dispute about the accretionary history of the major orogens, although significant second-order problems remain regarding internal terrane relationships. For example, in the Cordillera, a controversy still exists about the so-called Baja-BC problem. Paleomagnetic data suggest terranes displacements of 1,000–3,000 km parallel to the strike of the orogen (i.e., northwestward) but this is not supported by other structural data and remains an unresolved problem (Percival et al., 2004; see also discussion by Miall and Blakey, Chapter 1). There also remain general problems associated with low-angle normal faults, or detachment faults, such as those first mapped in the Basin and Range (Davis and Coney, 1979; Wernicke, 1985). Earthquake evidence for these faults is lacking, and it has proved difficult to accommodate this type of fault to mechanical theory (Axen, 2007). Many second-order problems, such as the regional relationships between tectonism, magmatism, and sedimentation, are now being tackled with the use of high-resolution geochronology, isotope geochemistry, and specialized geophysical applications, including re-processing of existing seismic data. Presented here are a few examples of some current problems and suggested solutions. This is not offered as a comprehensive overview of the ‘‘state-of-the-art’’ of sedimentary and tectonic analysis of basins in the United States and Canada, but as a flavor of the types of problems now being addressed. In the Appalachian orogen, controversy about the Pine Mountain Window, a complex faulted inlier of Grenville-age rocks straddling the Georgia–Alabama state line, represents a revisiting of the old problem about basement-uplift versus thin-skinned tectonics. Is the inlier the product of post-orogenic basement uplift along normal faults, or is it part of an allochthonous, overthrusted exotic (South American) terrane above the Alleghenian de´collement, which seismic data has been unable to adequately distinguish from Laurentian basement beneath (see Figures 12 and 13 in Chapter 1 for maps of the Alleghenian collisional orogen). The original COCORP seismic line across the inlier did not show the master Appalachian de´collement here, although it is very clear along the same line to the north, hence the basement-uplift interpretation. However, careful re-processing of this data by McBride et al. (2005) revealed significant reflections at the de´collement level, which they concluded had been obscured in the original data by poor acquisition conditions, relating partly to the rugged topography of the Pine Mountains. The master Appalachian de´collement can now be confidently traced beneath the entire orogen to the base of the crust at depths of nearly 40 km beneath the Atlantic coastal plain. Crustal shortening and overthrusting on this de´collement ranges from 100 to 400 km (Ettensohn, Chapter 4). At the north end of the Appalachian orogen in Newfoundland, the type area, in many ways, for platetectonics interpretations (as suggested above), it has long been known that William’s (1978) original assignment of central Newfoundland to a single ‘‘Dunnage terrane’’ required further study and elaboration. His original three subzones have undergone extensive remapping and analysis and it has now been shown that they comprise a complex of arc, oceanic, and continental fragments, revealing a complex history of rifting, spreading, and closure during the brief life of Iapetus Ocean (van Staal, 2005). Detailed studies reported from west-central Newfoundland (Zagorevski et al., 2006), including new mapping data, new, high-precision U/Pb dates and igneous geochemistry reveal the details of a west-dipping arc and microcontinent complex (Dashwoods terrane), the latter rifted from Laurentia during the Early Cambrian, and the entire complex welded back onto Laurentia during the Middle–Late Ordovician. Extension of what Zagorevski et al. (2006) call the Annieopsquotch accretionary tract and the indications of very Early Taconic west-dipping subduction into Quebec and New England are still the subject of controversy. However, Zagorevski et al. (2006) were able to confirm correlations along strike, that had been suggested earlier, with the Connemara Group in Ireland, and the Northern Belt of the Southern Uplands of Scotland. Ettensohn (Chapter 4) notes the occurrence of rifted Grenville fragments that separated from the US portion of the continent during the breakup of Rodinia. One of the interesting observations to have been made from this remapping is the rapidity with which major tectonic changes occurred, with the evolution of major backarc systems and the emplacement of the ensialic Buchans arc in Newfoundland in 5–7 million years. Zagorevski et al. (2006) also emphasized the importance of strike-slip deformation of the rocks in Newfoundland and in the British Isles, which complicates reconstructions of platetectonic scenarios. Studies continue on the complex, tectonized sedimentary rocks within the Appalachian orogen. For example, the Upper Ordovician Saint-Daniel Me´lange of southern Quebec, is part of the Internal Domain of the Humber zone, within the Taconic orogen. It occupies the space immediately to the southeast of the series of thrust slices described by Lavoie (Chapter 3, Figure 7). Field work by Schroetter et al. (2006) demonstrated that this me´lange was deposited in a deep-water, forearc setting, and incorporates chaotic olistostrome breccias generated by

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sedimentary slumping in a tectonically active environment. Schroetter et al. (2006) suggested comparisons with the lower Great Valley Group of California (Ingersoll, Chapter 11). During the Late Precambrian to Devonian, the western continental margin of Laurentia functioned as an extensional margin (Miall, Chapter 5). Miogeoclinal sedimentation was dominated by carbonate depositional systems. Between California and British Columbia, continental slope and base-of-slope depositional systems have been extensively disrupted by Cordilleran tectonism, and only in a few areas the deposits are well enough exposed to be amenable to detailed basin analysis. Stevens and Pelley (2006) described one such example, a base-of-slope submarine fan system of Middle Devonian age exposed in east-central California. The main exposures are in the Inyo Mountains, with additional exposures occurring as pendants in the Sierra Nevada to the west. Most of the major components of the fan, including channel, slope, and overbank deposits are preserved. Sediments were derived by erosion of the miogeocline to the east during a low stand of sea level, and transported across the shelf within erosional channels. The outer limit of Cordilleran deformation is defined by the trace of the fold-thrust belt. This shows major cratonward extensions, such as in the southern Northwest Territories, and intervening re-entrants where the fold-thrust belt wraps around a buttress in the continental crust. This curving trace of the fold-thrust belt is interpreted as a reflection of the thickness of the underlying continental crust, with most of the deformation taking place within areas of thinned crust characterized by high backarc heat flow during deformation (Hyndman et al., 2005). Recent work by Saltus and Hudson (2007) added an important element to this interpretation. They noted that three of the major westward-projecting buttresses, located in northeastern British Columbia, northcentral Yukon Territory, and beneath the British Mountains of northern Alaska, are characterized by significant aeromagnetic highs. Saltus and Hudson (2007) suggested that the buttresses represent areas of thick crust that have undergone large-scale mantle underplating. As noted in Chapter 5, mantle underplating beneath northeastern British Columbia is consistent with an earlier interpretation of this region of the Paleozoic continental margin as an upper-plate margin within a simple-shear model architecture (Lister et al., 1986). Detailed work continues to generate an even more precise picture of the structural evolution of the Sevier orogen in the Rocky Mountain states and its continuation northward into Canada. For example, Horton et al. (2004) and DeCelles and Coogan (2006) developed meticulous field documentation of the relationship between fault movement and sedimentation. Price and Sears (2000) developed a palinspastic reconstruction of the Rocky Mountains in southern Alberta, which provides constraints on the former western extent of the foreland basin. To the classic methods of cross-cutting relationships and the mapping of local clastic wedges can now be added the new techniques of direct dating of fault movement. Van der Pluijm et al. (2006) reported on the isotopic dating of clay-bearing fault gouge, which suggested the occurrence of two clusters of fault movement in the southern Canadian Rockies, at around 72 Ma and again at 52 Ma. Space does not permit a review of all the detailed stratigraphic and sedimentologic studies currently underway in the sedimentary basins of the United States and Canada. New chronostratigraphic techniques are providing more precise ages for important stratotypes and sequence boundaries, such as, for example, the carbon isotope studies of Sageman et al. (2006) and Gro¨cke et al. (2006). One of the main outcomes of such work is likely to be a much more tightly constrained chronostratigraphic framework for the Cretaceous rocks of the Western Interior Seaway (Miall et al., Chapter 9), from which an orbital time scale is gradually being constructed. Meanwhile, the very detailed regional sequence studies of individuals such as G. Plint and his group (e.g., Plint and Kreitner, 2007) continue to add to a high-resolution stratigraphic framework for the Cretaceous rocks of the Western Interior Seaway. Isopachs constructed from this framework can be interpreted to indicate local shifts in sediment supply and dispersal routes, and also clearly indicate both high-frequency tectonic and eustatic (glacioeustatic?) controls on accommodation.

ACKNOWLEDGMENTS This essay has benefited greatly from the different perspectives of Phil Allen and Peter DeCelles, who both provided many useful comments and suggestions. However, the author is deeply aware of the dangers in attempting a broad-brush synthesis of this type, and wishes to release these valued colleagues from any blame for remaining errors, inconsistencies, and omissions.

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Postscript: What have We Learned and Where Do We Go from Here?

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INDEX

Absaroka-I and II sequences in Paleozoic western craton margin, 205 Absaroka sequence of cratonic platform cover evolution (Late Mississippian to Early Jurassic), 46–49 distribution of, 15 sediments of, 297, 302 Absaroka strata, deposition of, 58 Absaroka volcanic field, 441, 442 detritus from, 439 Acadian basins, latest Ordovician to Middle Devonian, 93–94 Acadian orogeny, 36, 65 cause of, 6 termination of, 218–219 Acadian thrusts, reactivation of, 220 Accretion Alexander–Wrangellia–Peninsula terrane to North American margin, 370–371 history of, 23 Adirondack Massif, 107, 109, 110, 144 Aegean sea, subsidence of, 574 Aklak sequence, 562 frequency of reservoirs in Tertiary sandstones of, 568 Late Paleocene to Early Eocene, delta-plain deposits of, 560 Upper Paleocene to Lower Eocene, 567 Akpak sequence, 559, 562, 563 Alberta basin, 365, 372 Blairmore clastic wedge, 371, 372 Alberta, southern, Precambrian-lower Paleozoic section of, 187 Albian depositional environments, early, 560 end of, 554 Alexander terranes, 370, 385 Alexander–Wrangellia basement rocks, 371 Alleghanian cycles, 106 deformation, 139, 236 faulting, 232 Mountains, 155–156 tectonism, 162 onset of, 216 tectophase, 154, 161 Alleghanian orogeny, 16, 36, 49, 106, 109, 146–148 active, 302 collisional nature of, 152 development of, 55 Gondwana-Laurentia, 67 inception of, 149 onset of, 225, 226 Pennsylvanian-Permian, 12 product of, 107 tectonic influence of, 302 Allochthon, 69, 83 Allochthonous and Pericratonic terranes, 368 Allochthonous salt bodies, Gulf of Mexico basin, 510 Allochthonous terranes, 335, 365, 578 principal characteristics, Canadian Cordillera, 368 Allogenic process of sequence development in Western Interior basin, 349–357 Alluvial fan system and lacustrine margins, short links between, 439 Amalgamation of Quesnellia and Stikinia, final, 24 terrane, 23 Amarillo Uplift, 309 Amerasia basin, 454, 462 compression of Sverdrup basin, 467 geological setting, 454

rift activity related to initiation of, 462 tectonic event, 465 Amerasia ocean basin, 466 Amphistegina shale, Gulf of Mexico basin, 534 Anadarko Basin Bally, stratigraphic cross-section through, 314 Anadarko basins, 299, 308, 314, 323, 324, 325 in Absaroka subsequence, 48 Anahuac transgression, Miocene episodes, 534 Ancestral Rockies, and southern mid-continent, development of, 16 Ancestral Rockies basins, 248, 285 distribution of, 249 yoked, sedimentary fill of, 285 Ancestral Rocky Mountains, 299, 303, 325, 330 development of, 17 Uplifts of, 303 Anthracite basin, 156 Anticosti foreland basin, 84 Antitectonic model, 576 Antler belts, 399 Antler flysch, 399 Antler foreland basin, 250 depositional sequence, 399 exposures in central Nevada, 405–407 Mississippian clastic sequence, evolution, 404 siliciclastic deposits, 405 Antler foreland deposits, 399 Antler foreland strata, Late Mississippian, 405 Antler orogeny, 3 of California-Nevada, 578 creation of, 46 Devonian–Mississippian, 398–407 Antler proforeland basin, 397, 399, 400, 403 depositional sequence, 399 subsidence of edge of North America, 403–404 Apache basins, 249 Appalachian basins, 225, 233. see also Appalachian foreland basin, Canada Appalachian belt, of eastern United States, 309 Appalachian cyclothems, 302 Appalachian externides, 109 Appalachian foreland basin, 226 Appalachian foreland basin, Canada, 65, 576 hydrocarbon potential of, 94–95 in Kaskaskia subsequence, 47 Post-Taconian to Acadian basins, 82 Continental Eastern Canada, 84–90 Newfoundland, 83–84 paleogeographic reconstruction of, 90–92 regional geological setting, 66–67 sea level record in Lower to Middle Paleozoic Appalachians, 93–94 Taconian-deformed basins (The Humber Zone), 67 end rift–early drift, 69–73 Lower Paleozoic end-rift and passive margin, regional sea-level scenario for, 78 passive margin, 73–78 Taconian foreland basin, 78–82 tectonostratigraphic domains of, 67 Appalachian geology, subdivisions of, 574 Appalachian orogen(ies) clastic wedges developed due to, 9 deformed rocks of, 2 of eastern North America, 308 generation of, 1 in Newfoundland, 584 northern, terranes in, 214 series of, 6 terranes of, 578 Valley and Ridge province of, 16 Appalachian Plateau, 109

593

594

Appalachian shelf, 45 Aptian deepening, of Gulf of Mexico basin, 523 Arabia, convergence of, 235 Arctic margin of Franklinian basin, 206 plate-tectonic evolution of, 12 sedimentary evolution of, 16, 25 Arctic Ocean Canadian Beaufort-Mackenzie basin lies on, 551 multiphase opening of, 557 with sense of rotation, 556 Arctic platform, 346 Argo Salt, 489, 491 Arkansas Novaculite, 301 Arkoma basins, 310 diagrams for, 316 Aroostook-Perce´ Anticlinorium (APA), 85 Arthropods, remains of, 227 Artinskian Great Bear Cape Formation, 459 Asian geology, complex of, 578 Atkinson Point Formation, 567 Beaufort-Mackenzie basin, 567 Atlantic and Gulf margins, sedimentary evolution of, 25–26 Atlantic Canada, Maritimes basin of, 211–213 basement rocks and basinal overview of, 214 basinal events, overview of, 216–217 Gondwanan glaciation, 217–218 stratigraphic terminology, 215–216 terrane assembly, 214–215 geology and sedimentary research in, history of, 213–214 history, synopsis of, 233 key points, 233–234 magmatic history, 234–235 Late Devonian-Mississippian, 220 Fountain Lake Group, 222 Horton group, 220–222 Mid-Carboniferous: extensional basins and first phase of coal measures in, 225 lower part of Cumberland Group, 225–227 salt migration, 227–228 southwestern Maritimes basin, 227 Stellarton group, 227 Mid- to Late Devonian, 218 Acadian orogeny termination, 218–219 basins and batholiths, 219–220 Mississippian: global transgression and thermal subsidence, 222 Mabou Group and basin-margin facies, 224–225 Windsor Group, 222–224 Mississippian: tectonism and local basin filling, 222 Mississippian-Pennsylvanian unconformity, 225 modern analogue: Turkey and Eastern Mediterranean, 235–237 Pennsylvanian-Permian: second phase of coal measures, 228 Morien and Pictou Groups, 229–232 thermal subsidence, 228–229 Pennsylvanian to Permian sedimentation and tectonic events in, 232 climate change, 232–233 final tectonic events, 232 Permian eolian sandstones, 232 Permian to Mesozoic, 233 Atlantic coastal plain, 489 Atlantic margin basins of, 478 generalized cross-sections through, 481 and Gulf, sedimentary evolution of, 25–26 numerical model for flexural subsidence on, 478 petroleum resources, 499–500 uses of, 185 Atlantic Ocean best-fit closure of, 482 connection between Gulf of Mexico and, 486 crust extension, 494 plate-tectonic evolution of, 476 sea-floor spreading in, 477, 493 studies of paleomagnetic lineations in, 500 Atokan sequence, in Pennsylvanian, 256–257 Aulacogen, 300 Oklahoma, 309, 310, 324

Index

Austin depositional episode, Late Cretaceous, 527 Avalon Formation, 494 Avalon terrane, 214 Baca basin, 432 Back-arc basin, 251, 335 analysis, 290 development, 290 Backbulge basin, 403, 404 Baffin Bay, 496, 499 Bahama Platform, 482 deepwater channels of, 482 remote sensed image of, 483 Bahamas platform, 299, 315 Baja BC, 373–375 Baja British Columbia, 23 Baltimore Canyon trough, 486 continental margin off, 479 cross-section through, 489 location map of, 488 seaward edge of, 487 stratigraphy in, 488 Banff formation, 203 Bartlesville sandstone of Kansas and Oklahoma, 299, 311 model of, 312 Basal-Sauk strata, 43 Basaltic volcanic flows Sverdrup basin, Cenomanian, 466 Base level Sverdrup basin, Middle–Late Permian, 460 Basement control, Western Interior basin, 354–356 Basement deformation, phase of, 430 Basement rocks, and basinal overview of Maritimes basins of Atlantic Canada, 214 Gondwanan glaciation in, 217–218 stratigraphic terminology in, 215–216 terrane assembly in, 214–215 Basement structures, Gulf of Mexico basin, 509–510 Basinal deposits, 250 Basinal overview, and basement rocks of Maritimes basins of Atlantic Canada, 214 Gondwanan glaciation in, 217–218 stratigraphic terminology in, 215–216 terrane assembly in, 214–215 Basinal shales, 306 Basin analysis, numerical, 289 Basin-floor compressional fold belts, Gulf of Mexico basin, 510–511 Basin formation in Turkey, 235 Basin-Range Province, 430 Basins analysis, in North America, 573 basin tectonics, 573 stratigraphy and sedimentology, 579–581 (areas), covered in the book, 2 and batholiths, 219–220 distribution, and classification of Laramide, 430–432 intracratonic, North American Hudson Bay basin, 58–61 Illinois basin, 53–56 Williston basin, 56–58 Jurassic, tectonic setting of, 289 non-yoked, 285–287 Pennsylvanian-Permian, tectonic origins of, 283–284 post-Taconian, 83 post-Taconian to Acadian continental eastern Canada, 84–90 Newfoundland, 83–84 paleogeographic reconstruction of, 90–92 structure, Laramide, 435–437 tectonics, 573 thermal subsidence, and phase of coal measures in, 228–229 Triassic, tectonic setting of, 287 Uinta and Piceance, 439 yoked, 284–285 Basin subsidence, phases of, 493 Batholiths, in Mid- to Late Devonian, 219–220

595

Index

Bathonian sandstone, 279 Bay St. George basin, 222, 227 Beaufort foldbelt, 564 Beaufort-Mackenzie basin Canadian, geology of Late Cretaceous to, 551–553 geographic setting, 553 previous work, 553–554 delta-complexes, 562 economic geology of, 567–569 regional geologic setting for, 554 plate tectonic setting, 555–556 tectonic elements, 554–555 stratigraphy and sedimentology of, 556 sedimentology, 557–559 stratigraphic history, 559–564 stratigraphy, 556–557 structural traps, 568 structure and tectonics of, 564 basin-margin faults, 564 Beaufort foldbelt, 564–565 listric faults in, 566 timing of deformation, 566–567 Beaufort sea eastern, listric-growth faults in, 566 on southeast margin of Canada basin, 552 Belt basins, 249 Bengal basin, 313 Ben Nevis Formation, 494 Benue trough, 300 Berriasian depositional environments, late, 559 Besa River Formation, 203 Bighorn basins, map of, 442 Bighorn Mountains, Uplift of, 435 Bimini Bank, 484 Bioherms archeocyathid, of Lower Cambrian Age, 193 of Gays River Formation, 223 hosting Pb-Zn deposits, 224 Bjarni Formation, 496 Black Warrior basins, 147, 148, 299, 309 in Kaskaskia subsequence, 47 Blairmore clastic wedge in Alberta basin, 371, 372 Blake Plateau basin, 486–487 closely-spaced west–east cross sections through, 487 Blow River High, Beaufort Foldbelt, 564 Blow Trough, 564 Blueschist metamorphism Cache Creek rocks, 367–368, 375 Bluff sandstone, 282 Boothia Uplift, movements of, 582 Borup Fiord Formation, 456–457 Bowser basin, 368, 370, 371, 377–380, 382 strata, 379 Bowser basin, British Columbia, 368 provenance linkages, 380 regional shortening across, 372 stratigraphic record, 370 tectonostratigraphic foundations, 377–380 Bowser Lake Group, 378, 379 Bridge Creek Limestone, 356 Bridge River–Cadwallader–Methow terranes, 381 Bridge River terranes, 23, 370 British Columbia ancient continental margin of, 186 Kicking horse rim of, 207 northern Robson basin of, 193 in Sauk sequence, 193 Precambrian-Lower Paleozoic section of, 187 southern Cordilleran orogeny for, 205 during Jurassic, continental margin of, 206 Broad basement arch. see Okak Arch Broken-foreland basins, 396 Brothers Peak Formation, Sustut basin (piggyback), 380–381

Buckner Formation, 524 Gulf of Mexico basin, 517, 518 Burgess shale, 191 and Cathedral limestone, relationship between, 194 deposition of, 192 location, 193 Burgos basin, 532 Burin platform, 224 seismic profile across edge of, 224 Cabullona Group, Sonora, 432 Cache Creek chert, metamorphism, 377 Cache Creek complex, 18 Cache Creek oceanic crust, 367 Cache Creek–Quesnellia terrane boundary, 371 Cache Creek terranes, 367, 376 Cadwallader terranes, 370, 381 Calaveras Complex, 411 Caledonian/Ellesmerian orogenic belt, 453, 454, 467 California Great Valley basin, 574 California-Nevada, Antler orogen of, 578 Californian terranes, 386 Cambrian carbonates, deposition of, 249 Early paleogeographic reconstruction, 75 strata of, 193 Laurentian continental margin, 77 Middle, western craton margin during, 193 orogeny, evolution of, 67 rocks, stratigraphic cross-section of, 198 strata blanket of, 190 onlap pattern of, 188 of Rocky Mountains, 191 on western continental margin, 185 Cambrian-Ordovician shallow, paleogeographic evolution of, 65 Canaan Peak Formation, 434 Canada Appalachian Foreland basin of. see Appalachian Foreland basin, Canada basin. see also Beaufort-Mackenzie basin Beaufort sea on southeast margin of, 552 thermally subsiding margin of, 556 Cambrian-Ordovician section, 192 continental eastern in early-/late Silurian, 86–87 in Late Ordovician-early Silurian, 85–86 in Late Silurian-Middle Devonian, 87–88 eastern in Late Ludlovian, 89 St. Lawrence rift of, 4 northern Kaskaskia-I sequence, 199–202 Tippecanoe sequence, 195 during Pennsylvanian, paleogeographic map of, 229 post-Taconian to Acadian basins of, 78–82 regional geological setting of, 66–67 Rocky Mountains of, 182 sequential East-West cross-sections through, 5 western cross-section through craton and miogeocline of, 190 Devonian system of, 201 in Kaskaskia-II sequence, 203 Canadian Cordillera modern plate boundary, 373 morphogeological belts of, 365, 367 sedimentary basins, time-space relationship, 369 terranes, 365–375 Canadian craton, east-west restored sections of, 59 Canadian oil sands reserves, 357 Canadian petroleum province, 493 Canadian Rocky Mountains Paleozoic rocks of, 191 southern, and great basin in, 185–189

596

Canadian shield, 330, 335, 346 margin of, 43, 45 tectonic element of North American craton, 33–34 transgression and encroachment of, 58 Canyon Fiord Formation, 457, 458 Capitan Reef facies relationships in, 322 Guadalupian classic reconstruction of, 315 reconstructions of, 321 margin, development of, 319 in southern midcontinent and Ouachitas, 314 stratigraphy of, 297, 299 Caradocian Misty Point Formation, 83 Caradocian Saint-Victor Formation, 90 Caradocian Winterhouse Formation, 83 Carbonate mudstone, 439 Carbonate platform, 458, 486 Carbonate strata, deposition of, 55 Carboniferous Kayak Formation, Beaufort Foldbelt, 564 Carbon sequestration, 158 Caribbean region, plate-tectonic reconstructions for, 483 Carmel page-lower, in Jurassic rocks, 277 upper, in Jurassic rocks, 278–279 Carolina troughs, 486–487 continental margin off, 479 North America drift during Late Triassic, 479 sea-floor spreading, 477 Cassiar Platform, 189, 190, 193, 199, 203 Castlegate Sandstone, 338 of Campanian age, 354 of Mesaverde Group, 348 Cathedral limestone, and Burgess shale, relationship between, 194 Catskill siliciclastic wedge, in Kaskaskia subsequence, 47 Cenomanian–Turonian Kaskapau Formation, 353 Cenozoic depositional episodes in Gulf of Mexico basin, 528–529 Late, Continental glaciation during, 26–27 Phanerozoic tectonics and depositional history during, 251–252 Phanerozoic tectonics and depositional history of Colorado Plateau and southern Rocky Mountains, 251–252 stratigraphic succession, Gulf of Mexico basin, 519 Central basin platform, 299, 314–315, 322 Central Carboniferous platform, 223, 225, 226, 228, 231 Central Colorado Trough, 256, 257, 258, 262, 284, 285 Chaleurs Bay Synclinorium (CBS), 85 Cherokee group of middle Pennsylvanian age, 306 of Oklahoma, 315 Chert-lithic detritus, 439 Chinle basin, 288–289 tectonic setting of Triassic basins, 288–289 Chinle formation in Triassic rocks lower, 268–269 middle, 269–270 upper, 270–271 in upper fluvial system, 274 Chronology, regional plate-tectonic, 583 Chugach terranes, 372, 373 Cincinnati Arch, 45 in Kaskaskia subsequence, 47 Circle Cliffs Uplift, 432 in Utah, Campanian, 434 Clam Bank Formation, 83, 84 Claron Formation, 434 Clastic sediment supply, Gulf of Mexico basin during Early Cretaceous, 521–522 during Late Cretaceous, 526 Clastic wedges, 337, 576 tectonic cyclicity, Western Interior basin, 353 of Western Canada foreland basin, 335 Climate change, during Pennsylvanian to Permian, 232–233 interplay of, 442 setting, Gulf of Mexico basin, 539–541

Index

Coal-bed methane (CBM), 157 Coalification Bowser basin, 380 Coal measures first-phase (Mid-Carboniferous), of Maritmes basin of Atlantic Canada, 225–228 Pennsylvanian, exposures of, 213 second-phase (Pennsylvanian-Permian), of Maritmes basin of Atlantic Canada, 228–232 Coal moisture, measurement of, 582 Coal reserves Gulf of Mexico basin, 544 Sverdrup basin, 468 Western Interior basin, Upper Cretaceous rocks, 357 Coal window, in Sydney basin, 230 Coast Belt, Canadian Cordillera, 371–372 characteristics of, 367 crustal thickening, 371 Cobequid hills, Horton strata in, 222 Cockfield Formation, 531 Louisiana, 531 Coeval basins, 220 ‘‘Cold spots,’’ 40–41 Collision Intermontane terrane with North American margin, 368 Quesnellia and Stikinia, 367 Collisional retroforeland basins, US Cordillera, 397 Colorado Plateau deposition on, 432 lithosphere of, 430 Phanerozoic tectonics and depositional history, 251–252 Uplift of, 27 Colorado Plateau, and Southern Rocky Mountains Jurassic basins of, tectonic setting of, 289 Utah-Idaho trough, 289 Zuni sag, 289 Pennsylvanian-Jurassic sedimentary basins of, 245–246 location, 246 scope and organization, 247–249 stratigraphic interval, 246–247 Pennsylvanian-Middle Jurassic sequence stratigraphy of, 252–254 Jurassic, 271 Pennsylvanian, 254 Permian, 260–263 Triassic, 263 Pennsylvanian-Permian basins of, tectonic origins of, 283–284 Cordilleran basins, 287 non-yoked basins, 285–287 yoked basins, 284–285 Phanerozoic tectonics and depositional history of, 249 Cenozoic, 251–252 Cretaceous, 251 Early and Middle Paleozoic, 249–250 Jurassic, 251 Pennsylvanian-Permian, 250 Triassic, 250–251 Precambrian basement, and its control on Phanerozoic deposition in, 249 trends and lineaments, 249 younger Precambrian sedimentary basins, 249 Triassic basins of, tectonic setting of, 287 Chinle basin, 288–289 Eastern Cordilleran basin, 287 Moenkopi shelf, 287 pre-Shinarump paleovalleys and Shinarump deposits, 287–288 Colorado river system, integration of, 252 Colton Formation, 439 Columbia, southern British, Phanerozoic margin of, 13 Comparable forearc basins, 417 Conasauga Formation, 128 Conflicting models, Nevadan orogeny in California, 411 Connecticut Valley-Gaspe´ synclinorium (CVGS), 84, 85 Connemara Group, Ireland, 584 Consortium for Continental Reflection Profiling (COCORP), seismic work of, 575 Continental glaciation, during Late Cenozoic, 26–27 Continental margin evolution, Gulf of Mexico basin, 541–543 Contour-parallel ridges, 506

597

Index

Cordilleran basins, 257, 262, 287, 364, 365, 366 eastern, Triassic, tectonic setting of, 287 Pennsylvanian-Permian, tectonic origins of, 287 Cordilleran Foreland basin, Kaskaskia subsequence in, 47 Cordilleran margin, Proterozoic strata of, 13 Cordilleran orogen assembly of, 1 deformed rocks of, 2 fold-thrust belt of, 181 formation, and Pangea breakup plate-tectonic evolution during, 17 sedimentary evolution of Arctic, Atlantic and Gulf, and western margins and western interior, 23–26 phase of, 3 tectonism of, 17 terranes of, 578 Cordilleran orogeny for southern British Columbia, 205 tectonism of, 206 Cordilleran shelf basin, 45 Cordilleran superterranes, 365 Cordilleran tectonics, influence of, 559 Cordilleran terranes, 281, 373 Cornwall Arch, 467 Cotton Valley depositional episode, 519–520 Covey Hill Formation, 125 Craton. see also North American craton Canadian, 59 definition of, 32 of North America, evolution of, 581–582 tectonic differentiation of, 45 Cratonic basin, 306 Cratonic cover, controls on evolution of, 36 eustasy in, 36–38 extension and thermal re-equilibration, 38 intraplate stress, 38 magmatic controls, 41–43 mantle downwelling, 40–41 mantle insulation and supercontinent cycles, 40 subducting slabs, 39–40 Cratonic margins, as tectonic element of North American craton, 35–36. see also Craton margin, Paleozoic western Cratonic platform cover, Phanerozoic evolution of, 43 Absaroka sequence, 46–49 Kaskaskia sequence, 46 Sauk sequence, 43–44 Tejas sequence, 50–51 Tippecanoe sequence, 44–46 Zuni sequence, 49–50 tectonic element of North American craton, 34 Craton margin, Paleozoic western, 181–182 grand cycles, 198–199 historical background, 182–185 Kaskaskia-II sequence, 202–203 Great basin, 203 western Canada, 203 Laurentia, rifted margin of, 185 southern Canadian Rocky Mountains and great basin in, 185–189 Yukon and northwest territories in, 189–190 Lower to Upper Devonian (Kaskaskia-I sequence), 199 ancestral Uinta Uplift, 202 Great basin, 202 northern Canada, 199–202 Peace River Arch, 202 Middle Ordovician–Early Devonian (Tippecanoe sequence), 195 Great basin, 195–197 northern Canada, 195 Pennsylvanian–Permian (Absaroka-I and II sequences), 205 Sauk sequence and Cambrian-Ordovician shelf-to-basin transition, 190 Great basin: Nevada, Utah, Idaho, 193–195 kicking horse rim and burgess shale of southern Canadian Rocky Mountains, 191–193 northern British Columbia, 193 Yukon and northwest territories, 193 Triassic–Jurassic: termination of passive continental margin, 205–206 Crazy Hollow Formation, Eocene, 439

Cretaceous cycles of transgression in Canada, 342–343 Creek sandstone, 282 Crescent terranes, 372, 388 Cretaceous basin, 341 North America during plate tectonic setting of, 20 western margin of, 21 in Phanerozoic tectonics and depositional history, 251 rock springs Uplift of, 439 Cretaceous Interior Seaway. see Western Interior Seaway Cretaceous–Paleogene granite plutons, British Columbia, 371 Cretaceous-Paleogene stratigraphy, of Sverdrup basin, 457 Cretaceous rocks coal reserves of Western Interior basin, 357 stratigraphic cross-section, from Utah to Colorado, 338 Cretaceous strata, 39–40, 56, 58, 490 Crockerland, 25 Crustal extension, 479, 493, 500 Crustal shortening Insular Superterrane, Canadian Cordillera, 371, 372 Western Interior basin, 354 Crustal structure Gulf of Mexico basin, 507–509 Crustal thickening Coast belt, 371 Cumberland basin, 226, 228, 229, 230, 231, 232 Cumberland Group, in Mid-Carboniferous, 225–227 Curtis-Summerville sequence, Jurassic rocks in, 280–281, 284 Cutler Formation, 17 Cyclic sedimentation cyclothem concept for, 581 models of, 309 at shelf margin of southern midcontinent and Ouachitas, 306–308 Cyclothems concept of, for cyclic sedimentation, 581 in southern midcontinent and Ouachitas, 302–306 Dakota Sandstone, 357 Darriwilian-Caradocian Lourdes Formation, 83 Davis Strait, 498 Davis Strait Transform, 498–499 Deep basin, 460 Deepest-water cratonic basin, 302–303 Deep Panuke, 500 Deer Lake basin, 212, 222, 224, 227 Deformation Canadian Cordillera, 372–373 of Nanaimo Group rocks, 389 Queen Charlotte-Georgia-Tofino basins, 387–388 zones, Sverdrup basin, 453 Degerbo¨ls Formation, 460 Sverdrup basin, 460 Delaware basins, 300, 306, 315, 321, 322 Absaroka subsequence, 48 evaporite sedimentation in midland and, 320–323 modern stratigraphic synthesis of, 315 Delta migration, pattern of, 564 Denver basins, 284, 432 Absaroka subsequence, 48 Zuni subsequence, 50 Depositional environments, Sonomia terranes, 409 Depositional episodes, 580 Depositional episodes, Gulf of Mexico basin, 514–516 Cenozoic, 528–529 Early Cretaceous, 520–525 Early Pliocene–Quaternary, 536–538 Laramide, 529–530 Late Cretaceous, 525–528 Middle Cenozoic volcanism and related, 531–533 Middle Jurassic–Earliest Cretaceous, 516–520 Miocene, 533–536 Depositional framework Gulf of Mexico basin, 514–516 Depositional loading in Gulf of Mexico basin during Miocene, 535–536

598

Depositional sequence Antler foreland basin, 399 Gulf of Mexico basin, 514–516 Havallah basin, 408 Depositional systems, of Laramide, 437–443 Depressions Sverdrup basin, Early Carboniferous–Early Late Carboniferous, 455–457 Desmoinesian sequence, Pennsylvanian, 257–258 Detrital-zircon age, 439 probability, for McCoy Mountains Formation, 438 Detritus, chert-lithic, 439 Devils Claw Formation, Bowser basin, 379–380 Devils River Uplift, in Marathon area, 313 Devonian carbonates, deposition of, 249 Mid- to Late, 218 Acadian orogeny termination, 218–219 basins and batholiths, 219–220 North America during, plate-tectonic setting of, 10 strata, in selected parts of Atlantic Canada, 218 system, of western Canada, 201 Devonian (Kaskaskia-I sequence), lower to upper, 199 ancestral Uinta Uplift, 202 Great basin, 202 northern Canada, 199–202 Peace River Arch, 202 Devonian-Mississippian, Late Fountain Lake Group, 222 Horton group, 220–222 Devonian–Mississippian Antler orogeny, 398–407 Devonian Palliser Formation, 485 Alberta, 485 Devonian platform, 368 Dezadeash basins, 370, 372, 385 Dezadeash formation, 385 Diablo platform, 299, 314–315, 315 Diapiric evaporites, 479 Dinosaur fossils, 432 Dolomite, subordinate beds of, 197 Dolostone, 193, 197 Douglas Creek arch, 432, 440, 442 Drainage history, Gulf of Mexico basin, 538–539 Duchesne River Formation, 439 coarse-grained, 439 Dunkard basin, 154, 156, 157 Dunnage Zone, 69, 82 rocks, 85 Dynamic load, 577 Dynamic loading effect, 330 Eagle basin, 258 Absaroka subsequence, 48 Early Cretaceous depositional episodes in Gulf of Mexico basin, 520–525 paleogeographic evolution of Western Interior basin, 339–345 Early Pliocene–Quaternary depositional episodes in Gulf of Mexico basin, 536–538 Eastern shelf, 299, 322 East Texas basin, 526–527 Eclipse Trough, 499 fill, on Bylot Island, 477 Edwards Group, Gulf of Mexico basin, 523, 524 El Chanate Group, 432, 433 Ellesmerian orogeny, 56 El Tuli basin, 432 Emma Fiord Formation, 455 Sverdrup basin, 445 Emplacement Roberts Mountains allochthon, 399 Emsian Red Island Road Formation, 84 End rift-early drift correlation western Newfoundland-Que´bec in, 69–71 Que´bec reentrant in, 71–72 St. Lawrence promontory in, 69–71 in Taconian-deformed basins -Humber zone, 69–73

Index

Energy resources Gulf of Mexico basin, 543–544 Sverdrup basin, 467–469 Entrada sandstone, 279 upper, Jurassic rocks in, 278–279 Eocene climatic optimum (ECO), 442 Eocene deposition, Gulf of Mexico basin, 531–532 Eocene Upper Wilcox depositional episode, paleogeography, 530 Eolian, mudstone of, 262 Epeiric basin, 493 Epeirogeny, 577 Eskimo Lakes Arch, 554 Beaufort-Mackenzie basin, 554, 559 Eurekan orogeny, 25 Europe, variscides of, development of, 234 European basins, 237 European geology, complex of, 578 Eustasy, 292 components of, 42 role in control of cratonic cover evolution, 36–38 vs. tectonism, 93–94 Eutaw Formation, Gulf of Mexico basin, 526 Evaporative basins, 228 Evaporite sedimentation, 320-323 Exhumation Cache Creek rocks, 367–368, 377 Extrabasinal carbonate clasts, 271 Facies-cycle wedge, 580 Fall Line, 489 Fall-line, 486 Farallon Plate, subduction, 330 Fault-bounded basins, 227 Fernie Formation, 330, 339 Filling deep basin of Sverdrup basin during Triassic, 460–461 Fish river sequence, depositional facies of, 563 Flagstaff basin, 435, 439 green river formation of, 438 Flat slab, for Laramide deformation, 446 Flemish Pass, 500 Flexural foreland basins, 285 Flexural loading Western Interior basin, 354 of Western Interior basin during Late Cretaceous, 348 Florida coast line, 530 Fluvial dispersal systems, 439 Flysch, 311, 312 Forearc basins, 396, 397–398 post-Nevadan, 414–418 post-Sonoma, 409–411 Forebulge Western Interior basin, 336–337 Foreland basins, 251, 289, 309, 310, 399, 400, 403, 575 development, 279, 282 stratigraphic columns in, 317 subdivision of, 577 Foreland sedimentation, phases of, 400–403 Fort Crittenden deposition of, 433 formation, 432 Fort Norman line, 333 Fort Worth Basin, 309, 310 FortWorth basin, Absaroka subsequence, 48 Fossil Burgess shale, 191 dinosaur, 432 siliceous replacement of, 320 Fountain Lake Group, in Late Devonian-Mississippian, 222 Fracture zone Bahamas to Newfoundland, 482 Fragmentation of Sverdrup basin during Paleocene–Eocene, 466–467 Franciscan Complex accretionary growth, 414 accumulation, 417–418

599

Index

Franciscan subduction complex, 370 Franklinian basin, 16, 453 Arctic margin of, 206 Ellesmerian Orogen deformed, 453 Frazierian depositional model, 515–516 Fredericksburg episode Gulf of Mexico basin, 524 Fundy basin, 482 Georges Bank basin, 489 Galice Formation, 414 in Klamath, 414 Gambier basins, 370, 372, 385 Ganderia, accretion of, 88 Garden Hill (GH) discovery of, 95 position of, 71 Gas. see Oil and gas Gaspe´ basin, palinspastic restoration of, 84 Gaspe´ Belt, 84 of the Late Ordovician to Middle Devonian, 85 Lower Paleozoic belts, hydrocarbon potential of, 95 northern edge of, 86 post-Taconian, paleotectonic reconstruction of, 90 salinic unconformity and disturbance in, 88–90 Gaspe´ depositional basin in eastern Que´bec, palinspastic restoration of, 84 Gaspe´ peninsula northern segment of Gaspe´ belt in, 91 seismic section in, 89 Gays River Formation, bioherms of, 223 Generalized stratigraphic table for Western Interior basin, 338 Genetic stratigraphic sequence, Gulf of Mexico basin, 515 Geochemistry, isotope, use of, 584 Geochronology, 90 high-resolution, use of, 584 Geodynamic framework Western Interior basin, 335–338 Geological setting regional, of Canada, 66–67 Sverdrup basin, 453–455 Geology economic, of Beaufort-Mackenzie basin, 567–569 of Late Cretaceous to Canadian Beaufort-Mackenzie basin, 551–553 North American, 581 analysis of, 576 evolution of craton in, 581–582 nature of, 574 other effects of intraplate stress in, 583 sediment fluxes in, 582–583 and sedimentary research, importance of history of, 213–214 of Turkey, 236 Geometry, Laramide, 435 Geophysical applications, specialized, use of, 584 Georges Bank basin oceanic crust offshore, 489 reef system, 490 source-rock studies in, 491 Georgia–Alabama state line, 584 Georgia basin, 372, 375, 389 Provenance linkages, 390 Geosynclinal cycle, 577 Geosyncline, Kay’s elaboration of, 573 Gilmer Formation, Gulf of Mexico basin, 517 Glaciation continental, 26 Gondwanan, phase of, 225 phase of, 216 Glen Canyon Group, Jurassic rocks of lower, 273 upper, 274–275 Gondwana glaciation basement rocks and basinal overview, 217–218 phase of, 225 with Laurussia, collision of, 6, 12

in Pangean assembly, 215 West African segment of, 214 Grand Banks of Newfoundland, 475 basins of, 493–496 Grand Canyon basins, 249 Grand cycles, 198–199 Stephen-type cycles, 198 Sullivan-type cycles, 198 Granites, anorogenic, intrusion of, 43 Granite Wash, 202, 309 Gravina basin, 385 Gravina–Nutzotin basins, 372, 382 Gravina–Nutzotin–Dezadeash basins, 372 Gravina–Nutzotin–Dezadeash–Gambier basins, 371 Gravity tectonic structures, Gulf of Mexico basin, 510–511 Great Bahama Bank margin, 486 northwest–southeast cross section across northern part of, 484 sediments of northern, 485 Great basin Kaskaskia-II sequence, 203 Kaskaskia-I sequence, 202 Nevada, Utah and Idaho, 193–195 Tippecanoe sequence, 195–197 Great Bear Cape Formation, Artinskian, 459 Great Isaac well, 482 Great Valley forearc basins, 414–416 California, 371 Great Valley Group sedimentation, 414 subsidence and thermochronologic analyses, 416 turbidite sedimentation within, 414 Greenhorn Formation, 346 Bridge Creek Limestone, 356 limestone–shale cycles of, 348 Greenland, East, rifting of, 14 Green River basin distribution of facies in, 441 map of, 442 Green River Formation, 439 deposits of, 440–441 division of, 440 Uinta basin, 439 Grenville Front, 110 Grenville Mountains, peneplanation of, 26 Growth-fault, Gulf of Mexico basin, 510 Growth structure domains, Gulf of Mexico basin, 511–514 Guadalupe Capitan Reef, reconstructions of, 315, 321 Guadalupe Mountains, 299 Guadalupian sequence, Permian, 263 Gulf coast, structural elements of continental margin of, 298 Gulf margins, and Atlantic, sedimentary evolution of, 25–26 Gulf of Maine Platform, 489 Gulf of Mexico basin, 509, 514 climate and oceanography, 539–541 continental margin evolution, 541–543 crustal structure and origin, 507–509 depositional framework, 514–516 depositional history and paleogeography, 516–538 energy resources, 543–544 physiographic elements of, 506 sediment supply, 538–539 structural framework, 509–514 Gulf of St. Lawrence basin, 212, 213, 220–221, 228 sandstones, 232 succession, 229, 231 Gulf of St. Lawrence platform, 224 Gulf succession, formations of, 231 Gypsum, production of, 214 Halite, 439 Hamill Group, 186 Hare Fiord Formation, 458 Hartford basin, 482 Havallah basins, 248 evolution, 407–408

600

Hazen Trough, 190 Helikian age, belt-purcell supergroup of, 186 Herschel High, Beaufort Foldbelt, 564 Hibernia oilfield, 494 Hinge line, 487 Holbrook basins, 262, 285, 286, 292 Hopedale basin, 498 depth to basement and well locations in, 498 seismic cross-section, 499 synrift faults in, 497–498 Hornbrook Formation, 417 Horquilla limestone, 256 Horseshoe reef, 315 Horton Group, Late Devonian-Mississippian, 220–222 Horton strata, in Cobequid hills, 222 Hosston depositional sequence, 521–523 Hudson bay basin initiation of, 34 intracratonic, 58–61 schematic cross-sections of, 52 Hudson platform chronostratigraphic diagram of, 60 transgression and encroachment of, 58 Hudson Strait, 479 Humber Arm Allochthon, 69, 83 Humber terrane, 214 Humber Zone, 67 end rift–early drift, 69–73 hydrocarbon potential of Que´bec, 94 in western Newfoundland, 94–95 Lower Paleozoic end-rift and passive margin, regional sea-level scenario for, 78 passive margin, 73–78 Taconian foreland basin, 78–82 Hunic terranes, 309 Hydrocarbon-bearing wells location of, 553 Hydrocarbons, 157–158 potential of Appalachian basins of Canada, 94–95 production, 299 source of, 551 source rocks, 492 Hydrology, basin, 231, 439 Iapetus ocean closure of, 6 remnants of, 15 Ice sheets, development of, 217 Illinois basin, 45, 53–56, 145, 146, 302 evolution of, 34 initiation of, 44 Kaskaskia subsequence, 47 schematic cross-sections of, 52 Illinois cyclothems, 302 Illtyd Formation, 193 Imbrication, structural Nutzotin–Dezadeash–Gravina–Gambier basins, 384 Insular superterranes, 348, 365, 370–372, 380 Canadian Cordillera, 370–372 Intermontane superterrane, Canadian Cordillera, 366–370 sedimentary basins associated with, 375–382 Intermontane superterranes, 365, 367, 368 sedimentary basins associated with, 375–382 Intermontane terranes, 339, 340, 366–370 Intrabasinal carbonate clasts, 270, 271 Intracratonic basins, tectonic elements of North American craton, 34–35. see also Basins, intracratonic Intraplate stress concept of, 578 on North America geology, 583 role in control of cratonic cover evolution, 38 Inversion-flank basins, 432 Iran, Zagros region of, 235 Iroquois Formation, 489–490

Index

Isopachs of Early/Late Cretaceous strata in Rocky Mountain states, 336 Mississippian strata, 407 Isostasy, theory of, 573 Jeanne d’Arc basin commercial hydrocarbon production from, 500 correlation with Orphan basin, 495 crust extension and, 494 Lithoprobe seismic line, 496 petroleum exploration in, 500 stratigraphic section of, 493, 495 Jeanne d’Arc Formation Hibernia oilfield, 494 Joggins formation, 226 Juan de Fuca Plate, 373 subduction of, 388 Jurassic back arc basins, 248 continental margin of southern British Columbia during, 206 and Cretaceous strata identification of, 557 North America during, plate tectonic setting of, 19 North America during, western margin of, 21 paleogeographic evolution of Western Interior basin, 339 sedimentary basins, 249 volcanic rocks of, 438 Jurassic basins, tectonic setting of, 289 Utah-Idaho trough in, 289 Zuni sag in, 289 Jurassic Kingak Formation, Beaufort Foldbelt, 564 Jurassic rocks, 271–272, 486 Curtis Summerville in, 280–281 lower Glen canyon group of, 273 Morrison formation and younger Mesozoic events in, 282–283 page-lower carmel in, 277 in Paleotectonic settings of southwestern North America, 248 in Pennsylvanian-middle Jurassic sequence stratigraphy, 271 in Phanerozoic tectonics and depositional history, 251 temple cap sandstone in, 275–276 unconformity J-0 of, 272 unconformity J-1 of, 275 unconformity J-2 of, 276 unconformity J-3 of, 280 unconformity J-5 of, 281–282 unconformity J-sk of, 274 unconformity J-sup of, 277–278 upper carmel, entrada in, 278–279 upper Glen canyon group of, 274–275 Kaibab formations, 263 Kaibab Uplift in Zuni subsequence, 50 Kaiparowits basin, 432, 434 Kaiparowits formation, 434 Kankakee arch, 45 Kanosh shale, 197 Kansas Bartlesville sandstone in, 299, 311 cyclothems of, 302 Karst terrain, 491 Kaskaskia-II sequence, 202–203 great basin in, 203 western Canada in, 203 Kaskaskia-I sequence, 199 ancestral Uinta Uplift in, 202 great basin, 202 northern Canada in, 199–202 peace river arch in, 202 Kaskaskia sequence of cratonic platform cover evolution (mid-Early Devonian to Late Mississippian), 46 Kaskaskia strata diagrammatic cross-sections of, 37 subdivision of, 199 Kayak Formation, Carboniferous Beaufort Foldbelt, 564

601

Index

Kayenta formation and Navajo sandstone, 274 Kay’s elaboration of geosyncline, 573 Kiamichi Formation Fredericksburg episode, 524 Gulf of Mexico basin, 524 Kicking horse rim, 186, 191, 198 of British Columbia, 207 ‘Kicking horse rim’ and burgess shale of southern Canadian Rocky Mountains, 191–193 Kingak Formation, Jurassic Beaufort Foldbelt, 564 King Salmon thrust, 367 Kluane schist, Dezadeash basin, 385 Kootenay-Fernie clast wedge, 369 Kootenay Formation, 339, 369 Kugmallit sequence, depositional facies of, 563 Kugmallit Trough, 564 Laberge Group, Whitehorse trough, 375 Labrador basin. see Labrador shelf Labrador margin rifting on, 477, 479 Labrador Sea, 477 rifting, 496 spreading cessation, 499 Labrador Shelf evaporites in, 479 Hopedale basin, 499 petroleum resources, 500 Lacustrine margins, and alluvial fan system, short links between, 439 Laramide, 25, 430, 432, 435 correlation chart of, 434 cross sections, 437 deformed province of, 430 depositional systems, 437–443 fluvial systems of, 438 models for, 436 orogeny, 430 of Rocky Mountain states, 581 sedimentary basins, 429–430, 431 tectonics of orogeny of, 443–446 timing of deformation of, 432–435 Laramide basin(s) classifications, 433 distribution and classification of, 430–432 structure, 435–437 Laramide deformation, 430 flat slab for, 446 manifestation of, 432 patterns of, 443 Laramide depositional episodes, in Gulf of Mexico basin, 529–530 Laramide orogeny, 24, 330–331, 435 crustal blocks during, 50 Laramide province, 430, 432 distal part of, 446 drainage patterns for, 444 Laramide thrusts, vergence of, 435 Late Cenozoic, continental glaciation during, 26–27 Late Cretaceous depositional episodes in Gulf of Mexico basin, 525–528 paleogeographic evolution of Western Interior basin, 345–349 Late Cretaceous Redmond Formation, 498 Late Jurassic dextral transtensional basin, 372 Late Jurassic Nevadan orogeny, 411–414 conflicting models in California, 411 Oxfordian–Kimmeridgian Mariposa Formation, 414 Late Paleozoic basins, 482 Late Precambrian marginal basins, 248 Late Tournaisian basinal fill, 222 Laurentia, 4 Cambrian-Ordovician stratigraphy of, 198 Canadian segment of, 79 continental margin of, 66 formation of, 65

Paleozoic continental margin of, 68, 182 pre-drift sediments on western margin of, 188 Laurentia, rifted margin of, 185 southern Canadian Rocky Mountains and great basin in, 185–189 Yukon and northwest territories in, 189–190 Laurentian margin in Canada, duration of, 206 Laurentian plate, at the end of Precambrian, 4 Laurussia in Pangean assembly, 215 Leonardian, Permian basin during, 320 Leonardian sequence, lower and upper, in Permian, 262 Liard basin, subsidence of, 185 Liard Depression, 195 Liard line, 189, 190, 195, 333 Limestones Capitan, slope facies of, 320 Cathedral, and Burgess shale, relationship between, 194 shelf, deposition of, 304 Lindstro¨m Formation, Lopingian, 460 Listric faults Beaufort-Mackenzie basin, 566 Lithology, 231 for Chinle uranium ore deposits, 270 Lithoprobe line, 221 Lithosphere of Colorado Plateau, 430 mantle, 234 regional response of, 235 stretching of, 38 uses of, 579 Little Hat Top basins, 432 Llano Uplift, 299, 300 Local unconformities Tyaughton–Methow basin, 382 Logan’s line, 69 Long Point Group, 83, 84 Lopingian Lindstro¨m Formation, 460 Sverdrup basin, 460 Lower Cretaceous Fredricksburg depositional episode, paleogeography, 524 Lower Cretaceous Lower Hosston depositional episode, Paleogeography, 522 Lower Eocene Wasatch Formation, 439 short fluvial systems in, 439 Lower Miocene succession, Gulf of Mexico basin, 534 Lower Wilcox depositional episode, Gulf of Mexico, 529–530 Luning-Fencemaker belt, 410 Lyell Formation, 199 Maastrichtian deposits, 560 Maastrictian sedimentation, 251 Mabou group and basin-margin facies in Mississippian, 224–225 Mabou Group minibasins, formation of, 225 Maccrady Formation, 159 Mackenzie Bay sequence, 559 Mackenzie Mountains mount kindle formation of, 195 Mackenzie River, 25, 26 Mackenzie Trough, 553 MacLean Strait, 467 Magmatic control on evolution of cratonic cover, 41–43 history, of maritimes basins, 234–235 Magmatism Canadian Cordillera, 372–373 style of, 579 Mahogany oil shale bed, 441, 443 Malpeque basin, 220 Mannville Group, 344 architecture of, 354 oil reserves in, 357 stratigraphic characteristics, 342 Mantle downwelling, influence on cratonic cover evolution, 40–41 Mantle insulation and supercontinent cycles, dynamic topography related to, 40 Marathon area, Devils River Uplift in, 313 Marathon margin, in Absaroka subsequence, 48

602

Marathon Uplift, 299, 301, 308, 320–321 Marginal Uplift, of Sverdrup basin during Early–Middle Triassic, 461 during Mid-Aptian, 465–466 Marine basins, 509 in western Nevada, 251 Marine carbonates, 223 Marine transgression–regression cycle, 262, 263 Mariposa Formation, Oxfordian–Kimmeridgian, 414 Maritimes basin of Atlantic Canada, 211–213 climate curve for, 232 comprises, 212 deposition, end of, 233 Duckmantian to Permian strata of, 231 history, synopsis of key points in, 233–234 magmatic, 234–235 southwestern, in mid-Carboniferous, 227 tectono-stratigraphic overview of, 217 Markland Formation, 497 Massifs, 107–109 external, 110 internal, 110 Master faults, 493 Matador arch, 322 Maverick basin, 524, 538 Maximum flooding surface, Gulf of Mexico basin, 515 Mazatzal province, accretion of, 249 McCoy Mountains Detrital-zircon age probability for formation of, 438 McCoy Mountains Formation, 433, 439 Detrital-zircon age probability, 439 Mediterranean, eastern, and Turkey in modern analogue, 235–237 Megaslides, Gulf of Mexico basin, 541 Meguma Group formation of, 218 Georges Bank basin, 489 Meguma terrane, 218, 219, 227, 234 Memramcook Formation, 144 Mendocino triple junction, 373 Mesaverde Group Castlegate Sandstone, 348 San Rafael Uplift, 434 Mesozoic events, younger, in Jurassic rocks, 282–283 Mesozoic foreland basin, 277 Mesozoic stratigraphic succession, Gulf of Mexico basin, 517 Metamorphic age, for confirmation of Taconian age, 69 Metamorphism blueschist. see Blueschist metamorphism of Cache Creek chert, 377 of Omineca belt, Quesnel–Cache Creek–Stikinia interaction, 369 Methow terrane, 382 Michigan basin(s), 45 chronostratigraphic diagram of strata in, 53 evolution of, 34 formation of, 38 initiation of, 44 intracratonic, 51–53 in Kaskaskia subsequence, 47 Mid-Carboniferous: extensional basins and phase of coal measures, 225 lower part of Cumberland group in, 225–227 salt migration in, 227–228 southwestern maritimes basin in, 227 Stellarton group in, 227 Midcontinent–Permian basins, 298 Mid-Cretaceous unconformity, Gulf of Mexico basin, 524–525 Middle Cenozoic volcanism, Gulf of Mexico basin, 531–533 Eocene deposition, 531–532 Oligocene deposition, 532 Yegua strata, 531 Middle Eocene Crazy Hollow Formation, 439 Middle Jurassic–Earliest Cretaceous depositional episodes in Gulf of Mexico basin, 516–520 Middle Jurassic strata, 489 Middle–Late Triassic shales, Sverdrup basin, 469

Index

Middle Miocene sequence, Gulf of Mexico basin, 534–525 Midland basins, 299, 300, 306, 315, 322 Midway Formation, Gulf of Mexico basin, 529 Milankovitch cyclicity, Western Interior basin, 306, 356 Mineral resources, from Appalachian basin, 158–159 Miocene depositional episodes, Gulf of Mexico basin Anahuac transgression, 534 depositional loading and extension along Gulf shelf margin, 535 Lower Miocene succession, 534 Middle Miocene sequence, 534–535 Upper Miocene depositional episode, 535–536 Miogeocline, marine rocks of, 202 Mira terrane, 214 Mississippian carbonates, deposition of, 249 clastic sequence, Antler foreland basin, 404 early, Rhyodacite tuffs of, 310 global transgression and thermal subsidence, 222 Mabou group and basin-margin facies, 224–225 Windsor group in, 222–224 North America during, plate-tectonic setting of, 11 rocks of, 203 strata, isopach maps of, 407 time, Absaroka strata deposition during, 61 Mississippian-Pennsylvanian unconformity, 225 Mississippi river, 17, 44 Mississippi river valley evolution of, 583 Mississippi Salt basin, 509 Mississippi system, 306 Mississippi valley graben, 45 Missourian sequence in Pennsylvanian, 259 Modern plate boundary Canadian Cordillera, 373 Cenozoic Basins-Harbingers of, 387–390 Moenkopi formation, 271 Moenkopi sequence lower, in Triassic rocks, 265–266 upper, in Triassic rocks, 266–267 Moenkopi shelf in tectonic setting of Triassic basins, 287 Mohican Formation, Toarcian, 489 Mojave province, accretion of, 249 Moncton basins, 212, 220, 222, 231 Moose River, 108, 109 Morphogeological belts, Canadian Cordillera, 365 characteristics of, 367 Morrison formation, 330, 337, 339 in Jurassic rocks, 282–283 Morrowan sequence in Pennsylvanian, 254–255 Mount Kindle Formation, 195 Mt. Rodgers Formation, 124 Muav limestone, 250 Mudstone carbonate, 439 of eolian, 262 Murre Faults, 493 Nahcolite, 439 Nanaimo basin. see Georgia basin Nanaimo Group, 389 deposition, 390 Nansen Formation, 458 Sverdrup basin, 458, 459 Narragansett basin, 149, 225 Navajo sandstone, 278 Kayenta formation and, 274 Navarro Group, Gulf of Mexico basin Laramide crustal stress, 528 Maastrichtian stage, 527 Nemaha Uplift in Absaroka subsequence, 48 Neodymium geochemistry, uses of, 582 Neotectonic transpression, Great Valley, 416 Nevada eastern, pilot shale of, 202 great basin of, 182, 191 Pennsylvanian and Permian of, 205

603

Index

Nevadan orogeny cause of, 23 phases of, 251 Nevadan orogeny, Late Jurassic, 411–414 conflicting models in California, 411 Oxfordian–Kimmeridgian Mariposa Formation, 414 Newark basins, 477, 482 Newfoundland Appalachian orogen in, 584 clam bank belt in, 83–84 Dunnage zone of, 15 Fogo Seamounts, 491 fracture zone, 477, 481 Grand Banks of, 475, 493–496 in post-Taconian to Acadian basins, 83–84 salinic unconformity and orogeny in, 84 western, coast of, 218 western, in Humber zone, hydrocarbon potential of, 94–95 western, lower Paleozoic belts - Humber zone in, 94–95 western, Ordovician deposits in, 83 western, passive-margin succession in, 76 Newfoundland-Que´bec, western correlation of, 72–73 in end rift-early drift, correlation of, 72–73 in passive margin, correlation of, 76–78 in Taconian foreland basin, correlation of, 81–82 Nicola horst, British Columbia, 369 Nigeria modern Benue trough of, 300 Non-yoked basins, 285–287 Norphlet Formation, 516 Gulf of Mexico basin, 516 North America Appalachians of, 66 Atlantic margin of. see North American Atlantic margin basin analysis, 573 basin tectonics, 573 stratigraphy and sedimentology, 579–581 broad phases of Phanerozoic history of, 4–5 classical two-stage model for evolution of passive continental margin of central eastern, 480 continent, assembly of, 3 eastern margin during Early Cambrian, 8 geology, 581 evolution of craton in, 581–582 other effects of intraplate stress in, 583 sediment fluxes in, 582–583 plate tectonic setting during Early Cretaceous, 20 during Early Devonian, 10 during Early Permian, 12 during Early Silurian, 10 during Late Mississippian, 11 during Late Pennsylvanian, 11 during Middle Cambrian, 7 during Middle Jurassic, 19 during Middle Ordovician, 9 during Triassic, 18 of western margin, from Permian to end of Cretaceous, 21–22 remaining problems in, 584 rotational northwestward drift, 479 southwestern, Paleotectonic settings of, 248 western, Jurassic tectonics of, 251 western margin, evolution of plate tectonic setting (from Permian to end of Cretaceous) of, 21–22 breakup of Pangea, and formation of Cordilleran Orogen, 17–26 construction of Pangea, 6–16 development of southern mid-continent and ancestral Rockies, 16–17 westward drift of, relative to Panthalassa, 20 North America continent assembly of, 3 North American Atlantic margin basins and tectonic elements of, 474 eastern margin of, 476

Grand Banks of Newfoundland, 475 southern portion of, 477 North American Cordillera, 364. see also Canadian Cordillera North American craton cratonic cover, evolution of, controls on, 36 eustasy, 36–38 extension and thermal re-equilibration, 38 intraplate stress, 38 magmatic controls, 41–43 mantle downwelling, 40–41 mantle insulation and supercontinent cycles, 40 subducting slabs, 39–40 cratonic platform cover, Phanerozoic evolution of, 43 Absaroka sequence, 46–49 Kaskaskia sequence, 46 Sauk sequence, 43–44 Tejas sequence, 50–51 Tippecanoe sequence, 44–46 Zuni sequence, 49–50 definition of, 32 intracratonic basins, 51 Hudson Bay basin, 58–61 Illinois basin, 53–56 Michigan basin, 51–53 Williston basin, 56–58 margin of, 246 sedimentary cover of, Phanerozoic evolution of, 31–32 tectonic elements of, 32–33 Canadian shield, 33–34 cratonic margins, 35–36 cratonic platform, 34 intracratonic basins, 34–35 North American midcontinent, Pennsylvanian cyclothem sequence of, 305 North American plate, 298 Western Interior basin, 335 North Atlantic basins, 479 generalized stratigraphic columns for, 480 glacial activity in, 479 stratigraphy and structure, 479 Northern basin, 512 Northern shelf, 322 North Horn Formation, 439 North Sea basin, 494 petroleum source rocks in, 494 Nova Scotia Windsor group in Shubenacadie basin of, 223 Nutzotin basin, 384–385 Nutzotin–Dezadeash–Gravina–Gambier basins, 382–386 provenance linkages, 386 tectonostratigraphic foundations, 382–384 Ocean-basin affinities, 367 Oceanic crust offshore, 489 Oceanography, Gulf of Mexico basin, 539–541 Ochoco basin, 417 Oil and gas discoveries, 568 production in southern midcontinent and Ouachitas, 323–324 in Oklahoma, discovery of, 323 production, 94 shale, 440, 441, 443 Oil and gas reserves Rocky Mountain states, 357 Western Canada foreland basin, 357 Oil sands reserves, Canadian, 357 Okak Arch, 497–498 Okinawa Trough, 414 Oklahoma aulacogen, 324 Bartlesville sandstone in, 299, 311 Cherokee group of, 315 cyclothems of, 302 oil discovery in, 323 Oklahoma basins, 300 Old red continent, 8 Oligocene Frio depositional episode, Paleogeography, 533

604

Oligocene–Miocene detachment province, 514 Omineca Belt, 205 Oquirrh basin, 257 development of, 205 Ordovician age conglomerate of, 76 deposits, 250 in western Newfoundland, 83 North America during, plate tectonic setting of, 9 Ordovician-early Devonian (Tippecanoe sequence), middle, 195 great basin in, 195–197 northern Canada in, 195 Ordovician-early Silurian late, in continental eastern Canada, 85–86 Ordovician orogeny evolution of, 67 Ordovician period, 51 Ordovician sediments middle, isopach of, 300 Organic metamorphism measurement of, 582 Orogenic belt sevier, development of, 430 Orogeny Acadian, 65 Laramide, 430, 435 Laramide, tectonics of, 443–446 Salinian, 90 Taconian, 82 Orogrande basins, 262, 285, 292 Orphan basin, 494, 495, 500 correlation with Jeanne d’Arc basin, 495 development of, 493–494 oil exploration, 500 Otto Fiord Formation, 456–457 Ouachita basins, 310, 313, 316 Paleogeographic diagram for, 318 stratigraphic columns in, 317 Ouachita belt, oil and gas production from, 323 Ouachita deformation and sedimentation, 308 Ouachita-marathon orogeny, creation of, 46 Ouachita margin, in Absaroka subsequence, 48 Ouachitan orogen, deformed rocks of, 2 Ouachita orogeny, 49 development of, 55 Ouachitas, southern midcontinent, Permian basin and, 297–299 Capitan Reef in, 314 cyclic sedimentation at shelf margin of, 306–308 cyclothems in, 302–306 evaporite sedimentation in Delaware and midland basins in, 320–323 oil and gas production in, 323–324 Ouachita deformation and sedimentation in, 308 Paleozoic structural and stratigraphic setting of, 299–301 Outcrop regions, of Sonomia, 410 Outer Hinge Line (OHL), 564 Oxfordian–Kimmeridgian Mariposa Formation, 414 Nevadan orogeny, 414 Ozark Uplift, 299 Ozona platform, 322 Pacific Rim terrane, 372 Pahrump and Belt basins, 249 Pahrump basins, 249 Paleocurrent direction of Lower Creatceous gravel in Western Interior basin, 342 Paleoenvironmental analysis, 580 Paleogeographic evolution Gulf of Mexico basin, 516–538 Nutzotin and Wrangell Mountains basins, 384 Western Interior basin, 339–349 Paleogeographic orientation, US Cordillera, 398 Paleogeographic reconstruction of Western Interior basin, 344 Paleogeography, 84 Arctic, 25

Index

Eocene Upper Wilcox depositional episode, 530 Lower Cretaceous Fredricksburg depositional episode, 524 Lower Cretaceous Lower Hosston depositional episode, 522 of midcontinent, 304 Oligocene Frio depositional episode, 533 of Rocky Mountain, 429 of Sverdrup basin during Mesozoic, 461, 463, 464 Upper Cretaceous Austin depositional episode, 527 Upper Cretaceous Navarro depositional episode, 528 Upper Cretaceous Tuscaloosa-Woodbine depositional episode, 526 Upper Jurassic Smackover depositional episode, 522 Western Interior basin, 340, 343, 345, 347, 351 Paleogeology, sub-Triassic, 268 Paleolatitude trajectory, US Cordillera, 398 Paleozoic Appalachians lower to middle, in Canada, 93 Paleozoic carbonate platform, 214 Paleozoic strata, lower, development of, 39 Paleozoic tectonic provinces, western USA, 396 Palinspastic reconstruction, of southern California, 417 Palliser Formation, Devonian in Alberta, 485 Pangea assembly of, 308 breakup of, 4, 233 formation of, 67 Pangea breakup, and formation of Cordilleran orogen, 17 plate-tectonic evolution, 17 sedimentary evolution of Arctic margin, 25 Atlantic and Gulf margins, 25–26 western interior, 24–25 western margin, 23–24 Pangea, construction of, 6 plate-tectonic evolution, 6 sedimentary evolution of Arctic margin in, 16 eastern continental margin in, 15–16 interior and western continental margin in, 13–15 southern margin in, 16 Pangean assembly final stages of, 232 Laurussia in, 215 Pangean rifting, stages of, 235 Pangean terranes, 281 Pannonian basins, 236 Panthalassa Ocean, 191 initiation of, 186 subduction of, 17, 26 Paradox basins, 17, 257, 258–259, 261, 284, 285, 292 in Absaroka subsequence, 48 Passive margin St. Lawrence promontory in, 75–76 of Taconian-deformed basins, 73–78 Peace River Arch, 199, 202, 203, 583 Alberta, 583 Pearya terrane, 16 Pedregosa basins, 285, 286, 292 in Zuni subsequence, 50 Peninsula terrane, 370 Pennington Formation, 147 Pennsylvanian age Atokan sequence in, 256–257 cycles of, 299 Desmoinesian sequence in, 257–258 eastern Canada during, paleogeographic map of, 229 Middle Cherokee group of, 306 sediments of, 314 Mississippian-Pennsylvanian boundary of, 254 Missourian sequence in, 259 Morrowan sequence in, 254–255 North America during, plate-tectonic setting of, 11 in Pennsylvanian-middle Jurassic sequence stratigraphy, 254 unconformity 2 (C-4) of, 255–256 3 (C-5) of, 257

Index

8 (C-6) of, 258–259 10 of, 260 Virgilian sequence in, 260 Pennsylvanian basins, 254, 255–256, 283–284 Pennsylvanian cyclothem sequence, of North American mid-continent, 305 Pennsylvanian era, Absaroka strata deposition during, 61 Pennsylvanian lithofacies, upper, 302 Pennsylvanian-Middle Jurassic sequence stratigraphy, 252–254 Jurassic, 271 Pennsylvanian, 254 Permian, 260–263 Triassic, 263 Pennsylvanian-Permian Absaroka-I and II sequences, in Paleozoic western craton margin, 205 basins, summary map of, 255 basins, tectonic origins of, 283–284 Cordilleran basins in, 287 non-yoked basins in, 285–287 yoked basins in, 284–285 in paleotectonic settings of southwestern North America, 248 in Phanerozoic tectonics and depositional history, 250 phase of coal measures, 228 Morien and Pictou groups in, 229–232 thermal subsidence in, 228–229 rocks, cross-sections of, 253 sedimentation and tectonic events, 232 climate change in, 232–233 final tectonic events in, 232 Permian eolian sandstones in, 232 Pennsylvanian–Permian basins, 249 Pennsylvanian–Permian Pedregosa basins, 249 Pennsylvanian strata, 314–315 isopachs of, 313 marine and nonmarine deposition of, 53 Pericratonic terranes, 365 Principal characteristics, Canadian Cordillera, 368 Peri-Gondwanan Meguma terrane, 214 Permian age igneous rocks of, 56 sediments of, 205 basin, 284, 299, 300, 301, 303, 306, 310 and the Capitan Reef, 314–320 during Leonardian, 320 sedimentology of, 315 basin, southern midcontinent, Ouachitas and, 297–299 Capitan Reef in, 314 cyclic sedimentation at shelf margin of, 306–308 cyclothems in, 302–306 evaporite sedimentation in Delaware and midland basins in, 320–323 oil and gas production in, 323–324 Ouachita deformation and sedimentation in, 308 Paleozoic structural and stratigraphic setting of, 299–301 Guadalupian sequence and younger Permian rocks in, 263 isopachs, 321 lower and upper Leonardian sequence in, 262 North America during plate-tectonic setting of, 12 western margin of, 21 in Pennsylvanian-Middle Jurassic sequence stratigraphy, 260 Pennsylvanian-Permian boundary in, 260–261 rocks, younger, 263 unconformity P-sc of, 262 P-tw in, 262 Wolfcampian sequence in, 261–262 Permian eolian sandstones, in Pennsylvanian to Permian sedimentation, 232 Permian strata, 314–315 Permo-Triassic Sonoma orogeny, 408–409 Perry Formation, 144 Petrofacies analysis Relay Mountain Group deposition, 381–382 Petrography, 432

605

Petroleum geology, North American, 25 productivity, 299 in Sverdrup basin, 467–469 Petroleum exploration in Bahamas, 482 in Jeanne d’Arc basin, 500 in Scotian basin, 477, 500 Phanerozoic basins, 325 Phanerozoic evolution, US Cordillera, 418–419 Phanerozoic history of Hudson Bay basin, 58–61 of Illinois basin, 55–56 of Michigan basin, 51–53 of Williston basin, 56–58 Phanerozoic tectonics, and depositional history of Colorado plateau, 249 Cenozoic in, 251–252 Cretaceous in, 251 early and middle Paleozoic in, 249–250 Jurassic in, 251 Pennsylvanian-Permian in, 250 Triassic in, 250–251 Physiographic elements, of Gulf of Mexico basin, 506 Piceance basin, 439 Piceance Creek basins, 432 map of, 442 Piceance-Washakie basin, in Zuni subsequence, 50 Pictou groups capping redbeds of, 231 morien and, in Pennsylvanian-Permian, 229–232 Pilot shale, of eastern Nevada, 202 Pinal province, accretion of, 249 Piskahegan formation, 219 Plate boundary, modern Canadian Cordillera, 373 Cenozoic Basins-Harbingers of, 387–390 Plate motions Gulf of Mexico basin, 507 relative, Canadian Cordillera, 372–373 Plate-tectonic evolution, 17 during development of southern mid-continent and ancestral Rockies, 16–17 during Pangea breakup, and Cordilleran Orogen formation, 17–23 during Pangea construction, 6–13 Plate-tectonic regime, changes in, 4 Plate tectonic setting, 555–556 Platform, defined, 34 Pocahontas basin, 156, 157 Point Formation, Atkinson Beaufort-Mackenzie basin, 567 Polycyclic basinal history, 213 Porters Creek Formation, Gulf of Mexico basin, 529 Post-Nevadan forearc basins, 414–418 Post-Sonoma orogeny, 409–411 Post-Sonoma successor basin, 409–411 Pragian-Early Emsian, in reconstruction of post-Taconian basins, 92 Precambrian sedimentary basins, younger, 249 Precambrian tectonic elements Western Interior Seaway basement, 333 Pre-Chesterian age, distribution of strata of, 204 Pridolian/earliest Lochkovian, latest, in reconstruction of post-Taconian basins, 91 Princess Margaret Arch, 467 Proforeland basins, 396, 397 US Cordillera, 397 Prominent tectonic arch, 454 Prophet Trough, 202–203 Provenance characteristics sedimentary basins of Canadian Cordillera, 370 Provenance linkages Bowser basin, British Columbia, 380 Georgia basin, 390 Nutzotin–Dezadeash–Gravina–Gambier basins, 386 Queen Charlotte–Wrangell Mountains basins, 387 Tyaughton–Methow basins, 382 Whitehorse trough, 375–377 Purcell Anticlinorium, 336, 348

606

Quartz-lithic sandstones, 438, 439 Que´bec eastern, geological map of, 85 in Humber zone, hydrocarbon potential of, 94 lower Paleozoic belts - Humber zone in, 94 southern, with distribution of post-Taconian units, 86 Que´bec reentrant correlation of end-rift and passive-margin in, 74 in end rift-early drift, 71–72 in passive margin, 76 post-Taconian units in, 87 in Taconian foreland basin, 80–81 Queen Charlotte basins, 372, 386 Tertiary, 388–389 Queen Charlotte-Georgia-Tofino basins, 387–388 Queen Charlotte–Wrangell Mountains basins, 386–387 provenance linkages, 387 tectonostratigraphic foundations, 386 Queen Charlotte–Wrangell Mountains basins, 386–387 Queen City episode paleogeography, 530 Quesnel–Cache Creek–Stikine terrane amalgamation, 370 Quesnel–Cache Creek–Stikinia interaction, 369 Quesnellia, Canadian Cordillera, 367–369 Quesnell terrane, 8 Raanes Formation, 459 Sverdrup basin, 459 Rapid Depression, 564 Rapid sediment loading, Lower Wilcox episode, 529 Raton basins, 432, 435 Reciprocal sedimentation, concept of, 299 Recklaw transgression, Lower Wilcox episode, 530 Red alluvial deposits, 227 Red Desert Hanna basin, in Zuni subsequence, 50 Red Island Road Formation, 83–84 Redmond Formation, Late Cretaceous, 498 Red Mountain Formation, 134 Reelfoot-Illinois basin, in Kaskaskia subsequence, 47 Regional mapping, Western Interior basin, 354 Regional shortening across Bowser basin, 372 Rejuvenation Sverdrup basin during Early Cretaceous–Late Cretaceous, 465–466 Relative plate motions Canadian Cordillera, 372–373 Relay Mountain Group Tyaughton–Methow basin, 381–382 Remnant ocean basins, 397 US Cordillera, 397 Retroforeland basins, 397 Retrogradational slopes, Gulf of Mexico basin, 541 R1 event, 85–86 Rheic ocean, 310 basin, 301 closure of, 213 Rhyodacite tuffs, of early Mississippian age, 310 Rhyolite, 83 Richardson Trough, 190, 193, 195, 199 Richards sequence, 559 Richards strata, thickening of, 567 Rift basins, 233, 478, 481–482 formation of, 479 orientation and basement relationships of, 481 sedimentary succession in, 482 stratigraphy in, 488 Rifting, 477, 479 and drifting, 481 Ringbone Formation, 432 New Mexico, 432–433 Roadian Assistance Formation, Sverdrup basin, 460 Road River Formation, 193, 199 Roberts Mountains allochthon, 398–399 emplacement, 399 exposures in central Nevada, 405–407 Robson basin, of northern British Columbia, 193 Rock salt, production of, 214 Rockwood Formation, 134 Rocky Mountain foreland basin, 283

Index

Rocky Mountain province, 435 Rocky Mountain states isopachs of Early/Late Cretaceous strata in, 336 oil and gas reserves, Cretaceous, 357 Turonian sedimentation of, 348 Rodinia, breakup of, 3, 26, 297 Roho fault families, Gulf of Mexico basin, 510 Rome Formation, 127 Romer’s gap, 213 Rotational spreading Gulf of Mexico basin, 507–508 Rubio Peak Formation, 435 Rucker basin, 432 Sabine arch, 523–524 Sabinetown Formation, Gulf of Mexico basin, 529 Sable basin, 500 Sag basins, 235 Saglek basin, 498, 499 hydrocarbons in, 499 synrift faults in, 498 Salinic unconformity and orogeny, in western Newfoundland, 84 Salt diapirs, Gulf of Mexico basin, 510–511 Salt migration, in mid-Carboniferous, 227–228 Salt welds, Gulf of Mexico basin, 510 San Andreas transform fault, right-lateral, development of, 26 Sandstones compositions, fluvial, 439 Cotton Valley depositional episode, 519 Creek, 282 cross-bedded, 83 Entrada, 279 Permian eolian, 232 of Tansill formation, 322 Tapeats, fluvial deposits of, 250 temple cap, in Jurassic rocks, 275–276 un-named Permian, 232 Wingate, 271, 273 Sand Wash subbasin, 439 Sandy delta platform, 368 San Juan basin, 357, 432, 435 San Marcos arch, 523–524 San Rafael Uplift, 432, 439 North Horn Formation, 439 in Utah, Campanian, 434 Sarasota Arch, 509 Sauk sequence and Cambrian-Ordovician shelf-to-basin transition, 190 of cratonic platform cover evolution (Late Precambrian to Early Ordovician), 43–44 deposition, Cambro-Ordovician, 49, 51 development of, 189 Sauk strata diagrammatic cross-sections of, 37 thinning of, 56 Scotian basin petroleum exploration in, 477, 500 thermal evolution of upland region bordering, 491–492 Scotian shelf basin, 500 commercial hydrocarbon production from, 500 hydrocarbon source rocks, 492 Mesozoic and Cenozoic section of, 491 and slope, cross-section across, 493 stratigraphic table for, 492 Scurry County, correlation of electric logs across margin of shelf in, 308 Sea-floor spreading in central Atlantic Ocean, 477 Icelandic-type subaerial basalts formed during, 481, 483 during Late Cretaceous, 493 termination of, 497 Sea-level record, in Lower to Middle Paleozoic Appalachians in Canada, 93 scenario, for Lower Paleozoic end-rift and passive margin, 78 Sedimentary basins, 246 associated with Intermontane superterrane, 375–382 Canadian Cordillera, space-time relationships, 369

Index

Insular superterrane along inboard margin of, 382–386 along outboard margin of, 386–387 Younger Precambrian, 249 Sedimentary evolution of Alantic and Gulf margins, 25–26 of Arctic margin in, 16, 25 of eastern continental margin in, 15–16 of interior and western continental margin, 13–15 of southern margin, 16 of western interior, 24–25 of western margin, 23–24 Sedimentary uranium deposits, Gulf of Mexico basin, 544 Sedimentation cyclic cyclothem concept for, 581 models of, 309 at shelf margin of southern midcontinent and Ouachitas, 306–308 reciprocal, concept of, 299 syntectonic model of, 576 Taconian flysch, 84 Sediment fluxes, in North America geology, 582–583 Sedimentology, and stratigraphy of Beaufort-Mackenzie basin, 556 in North American basin analysis, 579–581 Sediment supply Gulf of Mexico basin, 538–539 Sverdrup basin, 462 Seismic line, 484–485 Seismic-refraction studies, 477 Seismic studies, 477 Atlantic margin, 481 Baltimore Canyon trough, 488–489 Bimini Bank, 484 diapiric evaporites, 479 Jurassic rocks, 486 transitional crust, 483 Selwyn basin, 189, 190, 195, 199, 201 Sevier Orogenic Belt, development of, 430 Shale lithified, 432 Mahogany oil shale bed, 441, 443 oil, 440 Shallow seas in Sverdrup basin latest Triassic–earliest Cretaceous, 461–465 Shawangunk Formation, 136 Shelf drowning Gulf of Mexico basin during Early Cretaceous, 523 Shinarump deposits, in tectonic setting of Triassic basins, 287–288 Shoaling cycle, 527 Shoestring sands, 306 Short fluvial systems, examples of, 439 Shubenacadie basin, 222 Sierras Pampeanas province, 443 Siliceous extrabasinal conglomerate, 267 Siliciclastic deposits, Antler foreland basin, 405 Siliciclastic rocks, 45 Silurian deposition of, 250 Late/Early, continental eastern Canada during, 86–87 North America during, plate-tectonic setting of, 10 reef systems, development of, 58 salinic Orogeny, expression of, 67 tectonism, for control of sedimentation, 91 Silurian-Devonian basin, of Canadian Appalachians, 67 Silurian-Middle Devonian, continental eastern Canada during, 87–88 Skeena basin, 372 Skeena deformation, 372 Skeena fold belt regional shortening across Bowser basin produced, 372 Slats Creek Formation, 193 Slide Mountain-Quesnellia-Cache Creek-Stikine terrane, 367 amalgamation, 367 Slide Mountain terrane, 367 Sligo Formation, 523 Gulf of Mexico basin, 523 Slope Diapiric Province, 491

607

Sloss sequences, 14 development of, 37, 43 Smackover Formation, Gulf of Mexico basin, 517 Smoking hill formation, 564 sequence, 569 Smoky Group, 356 SNORCLE reflection seismic profile, Canadian Cordillera, 371 Sonoma(n) orogeny(ies), 15, 23, 251, 287 formation of, 290 Permo-Triassic, 184, 408–409 of Nevada, 17 Sonoma orogenic belts, 399 Sonsela sandstone, base of, 269 Southern basins, 430 Southern midcontinent, Permian basin, Ouachitas and, 297–299 Capitan Reef in, 314 cyclic sedimentation at shelf margin of, 306–308 cyclothems in, 302–306 evaporite sedimentation in Delaware and midland basins in, 320–323 oil and gas production in, 323–324 Ouachita deformation and sedimentation in, 308 Paleozoic structural and stratigraphic setting of, 299–301 Southern Rocky Mountains and Colorado plateau Pennsylvanian-Jurassic sedimentary basins of, 245–246 location, 246 scope and organization, 247–249 stratigraphic interval, 246–247 Pennsylvanian-Middle Jurassic sequence stratigraphy of, 252–254 Jurassic, 271 Pennsylvanian, 254 Permian, 260–263 Triassic, 263 Phanerozoic tectonics and depositional history of, 249 Cenozoic, 251–252 Cretaceous, 251 Early and Middle Paleozoic, 249–250 Jurassic, 251 Pennsylvanian-Permian, 250 Triassic, 250–251 Precambrian basement and its control on Phanerozoic deposition in, 249 trends and lineaments, 249 younger Precambrian sedimentary basins, 249 tectonic origins of Pennsylvanian-Permian basins of, 283–284 Cordilleran basins, 287 non-yoked basins, 285–287 yoked basins, 284–285 tectonic setting of Jurassic basins of, 289 Utah-Idaho trough, 289 Zuni sag, 289 tectonic setting of Triassic basins of, 287 Southern segment, basins of, 482–486 South Florida–Bahamas basin, 482–486 Sparta depositional episode, 530 Spray River Group, 205 Spreading, Gulf of Mexico basin, 507–509 Springhill Mines Formation, 227 St. Anthony basin, 212, 222, 223, 227, 228 St. Lawrence basin, southern Gulf of, 221 St. Lawrence Platform, 69 St. Lawrence promontory correlation of end-rift and passive-margin in, 74 in end rift-early drift, 69–71 in passive margin, 75–76 in Taconian foreland basin, 79–80 St. Lawrence rift, of eastern Canada, 4 St. Lawrence succession, Gulf of, 229, 231 St. Mary’s basin, 212, 222 Stellarton basin, 212, 227, 228, 232 coals, 236 formation of, 228 Stellarton group, in Mid-Carboniferous, 227 Stephen-type cycles, 198 Stikine terrane, 8 Stikinia, Canadian Cordillera, 367–368 Stratigraphic cross-section Western Canada foreland basin, 332

608

Stratigraphic sequence development Western Interior basin, allogenic mechanisms, 349–357 Stratigraphic table for Jeanne d’Arc basin, 495 for Scotian basin, 492 Stratigraphy and sedimentology of Beaufort-Mackenzie basin, 556 in North American basin analysis, 579–581 Stress Acadian orogeny, 46 intraplate, 38 other effects of, 583 Structural cross-section Western Canada foreland basin, 332 Structural evolution, of Taiwan collision zone, 414 Structural framework Gulf of Mexico basin, 509–514 Structural growth history, Gulf of Mexico basin, 514 Structural imbrication Nutzotin–Dezadeash–Gravina–Gambier basins, 384 Stump Formation, 330 Subducting slabs, dynamic topography related to, 39–40 Submarine canyons, Gulf of Mexico basin, 541 Subsidence of edge of North America to from Antler proforeland basin, 403–404 Great Valley Group, 416 Gulf of Mexico basin, 509 phases of basin, 493 Sverdrup basin during Mesozoic, 462, 465–466 Successor basins post-Sonoma, 409–411 US Cordillera, 397–398 Sullivan-type cycle, 198 Superterranes, 365 Sustut basin, 372, 380–381 Suwanee Channel, 487 Suwannee Strait, 486 Sverdrup basin Carboniferous and Permian units stratigraphic sequence, 455 deformation zones, 453 development of, 25 economic geology, 467–469 geological setting, 453–455 sediment source, 454 tectonic episodes, 467 tectonic evolution, phases, 455–467 Sweetgrass Arch, 337 Sweetwater trough, 285 Sydney basin, 212, 213, 219, 220, 224, 225, 228, 229, 230, 232 model for coal window in, 230 offshore, 220 stratigraphic column for, 219 Sydney mines, formation of, 229 Synrift faults, 497 Syntectonic model, of sedimentation, 576 Taconian age, metamorphic age for confirmation of, 69 Taconian-deformed basins (Humber Zone), 67 end rift–early drift, 69–73 Lower Paleozoic end-rift and passive margin, regional sea-level scenario for, 78 passive margin, 73–78 Taconian foreland basin, 78–82 Taconian flysch sedimentation, 84 Taconian foreland basin correlation western Newfoundland-Que´bec in, 81–82 Que´bec reentrant in, 80–81 St. Lawrence promontory in, 79–80 Taconian orogeny, 6, 36 expression of, 82 of northern Appalachians, 576 Taglu sequence delta deposits in, 561 depositional facies of, 563 Taglu strata, thickening of, 567 Taiwan collision zone, structural evolution, 414 Taku terrane, 371, 384, 385

Index

Tanana terranes, 370 Tango Creek Formation Sustut basin, 380 Tanquary High, 452, 453, 454, 459 Tansill formation, sandstone of, 322 Taos Trough, 284, 286 Tapeats sandstone, fluvial deposits of, 250 Tarahumara Formation, 433 Sonora, 433 Tectonic cyclicity, Western Interior basin, 352–354 Tectonic depressions, 456–457 Tectonic elements Canadian shield in, 33–34 cratonic margins in, 35–36 cratonic platform in, 34 intracratonic basins in, 34–35 of North American craton, 32–33 Western Interior basin, 334 Western Interior Seaway basement, Precambrian, 333 Tectonic episodes, Sverdrup basin, 467 Tectonic evolution Sverdrup basin, phases, 455–467 Tectonic provinces, western USA Paleozoic, 396 Tectonic quiescence, Sverdrup basin Late Carboniferous–Early Permian, 457–459 Late Cretaceous, 466 Tectonic realignment in Gulf of Mexico during Lower Wilcox episode, 530 Tectonic subsidence, components of, 42 Tectonic thickening of Omineca belt, Quesnel–Cache Creek–Stikinia interaction, 369 Tectonism, 477, 483, 486, 499 convergent, 3 of Cordilleran orogeny, 17, 206 episode of, 24 evidence of, 227, 443 interplay of, 442 and local basin filling, 222 of northern Cordilleran Margin, 189 role of, 86 thin-skinned, 576 Tectonostratigraphic foundations Bowser basin, British Columbia, 377–380 Nutzotin–Dezadeash–Gravina–Gambier basins, 382–384 Queen Charlotte–Wrangell Mountains basins, 386 Tyaughton–Methow basins, 381–382 Whitehorse trough, 375 Tejas sequence (Late Paleocene to present), of evolution of cratonic platform cover, 50–51 Temple cap sandstone, in Jurassic rocks, 275–276 Terminal deepening, Gulf of Mexico basin, 523 Termination, Western Interior Seaway, 330–332 Terrane allochthonous. see Allochthonous terranes analysis, philosophy and methodology of, 578 assembly, in basement rocks and basinal overview, 214–215 Canadian Cordillera, 365–375 defined, 364 pericratonic. see Pericratonic terranes types of, 364 Wrangell, composition of, 23 Tertiary Queen Charlotte basin, 388–389 Tertiary strata, listric-growth faults in, 566 Tesnus formation, 301 Texarkana platform, 314 Texas north central, reciprocal sedimentation in, 310 Permian basin in, 323 southwest, structural elements in, 301 Thermal re-equilibration, and extension influence on cratonic cover evolution, 38 Thermal subsidence basins, and phase of coal measures in, 228–229 Thermochronologic analyses, Great Valley Group, 416 Thermochronology, 435 Tibbit Hill Formation, 69

609

Index

Tippecanoe sequence, 195 of cratonic platform cover evolution (Middle Ordovician to Early Devonian), 44–46 deposition of, 55 great basin in, 195–197 northern Canada in, 195 Tippecanoe strata characteristics of, 60 diagrammatic cross-sections of, 37 Tobosa basin, 300, 301 Tofino basin, 372, 389–390 Topographic basin, 275 Topography, 430 basin, initiation of, 314 depositional, 299 dynamic, 577 amplitude of, 39 development of, 43 influence of, 49 related to mantle insulation and supercontinent cycles, 40 Transcontinental arch, 45, 249–250, 285, 583 transgression and encroachment of, 58 Transcurrent-fault basins, 397 US Cordillera, 397 Transgression Cretaceous cycles in Canada, 342–343 of Sverdrup basin during Mid–Late Triassic, 461 Transylvanian basins, 236 T2-R2 events, 86–87 T2-R3 events, 87–88 Triassic basins, tectonic setting of, 287 Chinle basin in, 288–289 eastern Cordilleran basin in, 287 Moenkopi shelf in, 287 pre-Shinarump paleovalleys and Shinarump deposits in, 287–288 North America during plate tectonic setting of, 18 western margin of, 21 rocks, 263–264 lower Chinle formation in, 268–269 lower Moenkopi sequence in, 265–266 middle Chinle formation in, 269–270 in paleotectonic settings of southwestern North America, 248 in Pennsylvanian-middle Jurassic sequence stratigraphy, 263 in Phanerozoic tectonics and depositional history, 250–251 unconformity Tr-cr of, 270 unconformity Tr-lm of, 266 unconformity Tr-1 of, 264 unconformity Tr-3 of, 267 unconformity Tr-sm of, 269 upper Chinle formation in, 270–271 upper Moenkopi sequence in, 266–267 Triassic basins, tectonic setting of, 287–289 Triassic-Jurassic, termination of passive continental margin in, 205–206 Triassic-Jurassic stratigraphy, of Sverdrup basin, 457 Triassic sedimentary basins, 249 Trold Fiord Formation, Sverdrup basin, 460 Tuktoyaktuk peninsula, normal faults in Eskimo lakes fault zone on, 565 Turkey, and eastern Mediterranean in modern analogue, 235–237 Turkey, basin formation in, 235 Turonian sedimentation, of Rocky Mountain states, 348 Tuscaloosa Formation, 525 Gulf of Mexico basin, 525 Tuscaloosa-Woodbine depositional episode, Late Cretaceous, 525–527 clastic sediment supply, 526 paleogeography, 526 Tyaughton–Methow basins, 370, 381–382 correlation with Bowser basin, 370 provenance linkages, 382 tectonostratigraphic foundations, 381–382 Tyaughton–Methow trough, 370 sedimentation, 372 Uinta basins, 432, 439 map of, 442

Uinta Formation, 442 Uinta trough, 249 basins, 249 Uinta Uplift, 439 Uncompahgre Uplift, in Absaroka subsequence, 48 United States. see also North America pre-drift sediments on western margin of Laurentia in, 188 subsidence plot for precambrian to Devonian strata of, 187 Uplift of Sverdrup basin during Paleocene–Eocene, 466–467 Upper Cretaceous Austin depositional episode, paleogeography, 527 Upper Cretaceous Navarro depositional episode, paleogeography, 528 Upper Cretaceous Tuscaloosa-Woodbine depositional episode, paleogeography, 526 Upper Jurassic Smackover depositional episode, paleogeography, 522 Upper Jurassic strata, 489 Upper Miocene depositional episode, Gulf of Mexico basin, 535–536 Upper Taylor episode deposition, Late Cretaceous, 527 UpperTriassic back arc basins, 248 Uranium deposits, Gulf of Mexico basin, 544 US Cordillera Devonian–Mississippian Antler orogeny, 398–407 Havallah basin, 407–408 Late Jurassic Nevadan orogeny, 411–414 Permo-Triassic Sonoma orogeny, 408–409 Phanerozoic evolution, 418–419 post-Nevadan forearc basins, 414–418 post-Sonoma successor basin and forearcs, 409–411 Utah miogeocline of, 197 San Rafael and Circle Cliffs Uplifts, 434 Utah-Idaho trough, 251, 271, 275, 277, 279, 289, 292 Val Verde foreland basin, 312 van Hauen Formation, 460 Sverdrup basin, 460 Variscides of Europe, development of, 234 Vibroseis technique, 575 Vicksburg Group, 532 Vinini Formation, 405 Virgilian sequence, in Pennsylvanian, 260 Volcanic-lithic detritus, 439 Volcanic rocks, of Jurassic age, 438 Volcanism, 235 influence of, 430 Volcanism, Gulf of Mexico basin Late Cretaceous, 509–510 Middle Cenozoic, 531–533 Wamsutta Formation, 144 Wasatch formation, 441 clastic sedimentary wedges, 441 Desertion Point Tongue, 441 Lower Eocene, 439 sedimentary wedges of, 441 Wasatch line, 248, 249–251, 254, 264, 267, 287, 289 Washita depositional episode, 524 Wash Limestone Member, 191 Waterfowl Formation, 198 Weches Formation, Gulf of Mexico basin, 530 Wedge-top basins, 396, 397 US Cordillera, 397 West Alberta Arch, 189, 195, 199, 202 Western Canada foreland basin clastic wedges of, 335 oil and gas reserves, 357 structural/stratigraphic cross-section, 332 Western Canada Sedimentary basin, stratigraphic cross-section through, 15 Western interior, sedimentary evolution of, 24–25 Western interior basin, 435 coal reserves, 357 evolving configuration, 337 generalized stratigraphic table for, 338 geodynamic framework, 335–338 geological theory, 332 initiation, 330

610

Milankovitch cyclicity, 356 paleogeographic evolution, 339–349 tectonic elements, 334 Western interior foreland basin, 23, 24–25 Western interior Seaway basement, Precambrian tectonic elements, 333 extension of, 331 termination, 330–332 Western margin, sedimentary evolution of, 23–24 Wetland basins, 228 Whitehorse trough, 367, 375–377 provenance linkages, 375–377 tectonostratigraphic foundations, 375 Williston basin, 45, 193 evolution of, 34 formation of, 38 intracratonic, 56–58 in Kaskaskia subsequence, 47 schematic cross-sections of, 52 Wilson Cycle, 576, 578 Wind River Uplift Laramide sediment during Middle Eocene, 440 Wilkins Peak Member deposition, 443 in Zuni subsequence, 50 Windsor group base of, 228 in Mississippian, 222–224 in Shubenacadie basin of Nova Scotia, 223 Wingate sandstone, 271, 273 Wisconsin arch, 45 Wolfcampian sequence, in Permian, 261–262 Woodruff Formation, 494

Index

Wrangellia terrane, 370, 385 Wrangell Mountains basin, 386 Wrangell Mountains-Queen Charlotte basins, 372 Wrangell terrane, composition of, 23 Yakutat terrane, 390 Yarmouth Arch, 489 Yavapai province, accretion of, 249 Yegua delta, 531 Yegua Group, Gulf of Mexico basin in Texas, 531 Yegua strata, Gulf of Mexico basin, 531 Yoho national park Cambrian-Ordovician section of, 192 Rocky Mountains of, 182 Yoked basins, 284–285 in tectonic origins of Pennsylvanian-Permian basins, 284–285 Younger Precambrian sedimentary basins, 249 Yucatan terranes, 309 Yukatat terrane, 373 Yukon northern, Mountainous areas of, 182 and northwest territories, 189–190, 193 Yukon-Tanana terrane, 370 rocks, correlation, 371 Zuni sag, in tectonic setting of Jurassic basins, 289 Zuni sequence of cratonic platform cover evolution (Middle Jurassic to Early Paleocene), 49–50 deposition of, 56