Gulf of Mexico Origin, Waters, and Biota: Volume 3, Geology [1 ed.] 9781603442930, 9781603442909

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Gulf of Mexico Origin, Waters, and Biota Volume 3, Geology

John W. Tunnell Jr., General Editor

With gratitude to the U.S. Geological Survey for supporting this book’s coeditors as well as many of its contributors, and for funding the color reproduction.

•

Gulf of Mexico Origin, Waters, and Biota Volume 3, Geology

␥␥␥ Edited by

Noreen A. Buster and Charles W. Holmes

Texas A&M University Press College Station

Copyright © 2011 by Texas A&M University Press Manufactured in the United States of America All rights reserved First edition This paper meets the requirements of ANSI / NISO Z39.48–1992 (Permanence of Paper). Binding materials have been chosen for durability.

Library of Congress Cataloging-in-Publication Data Gulf of Mexico origin, waters, and biota / [edited by John W. Tunnell Jr., Darryl L. Felder, and Sylvia A. Earle]—1st ed. v. cm.—(Harte Research Institute for Gulf of Mexico Studies series) Includes indexes. Taken from the Harte Research Institute for Gulf of Mexico Studies website: Gulf of Mexico origin, waters, and biota, is an updated and enlarged version of the Gulf of Mexico: its origin, waters, and marine life, first published by U.S. Fish and Wildlife Service in Fishery bulletin, v. 89, 1954. Contents: V. 1. Biodiversity / edited by Darryl L. Felder and David K. Camp ISBN-13: 978-1-60344-094-3 (cloth : alk. paper) ISBN-10: 1-60344-094-1 (cloth : alk. paper) 1. Mexico, Gulf of. 2. Marine biology—Mexico, Gulf of. 3. Geology—Mexico, Gulf of. 4. Oceanography—Mexico, Gulf of. I. Tunnell, John Wesley II. Felder, Darryl L. III. Earle, Sylvia A., 1935–. IV. Camp, David K. V. Series. QH92.3.G834 2009 578.77'364—dc22 2008025312 Vol. 2, Ocean and Coastal Economy ISBN-13: 978-1-60344-086-8 ISBN-10: 1-60344-086-0 Vol. 3, Geology ISBN-13: 978-1-60344-290-9 ISBN-10: 1-60344-290-1

To all the scientists who have devoted their unending talents toward a more thorough understanding of and appreciation for the Gulf of Mexico . . . past, present, and future.

Contents

Dedication

v

Foreword: Fifty-Year Update of Bulletin 89 John W. Tunnell Jr., Darryl L. Felder, and Sylvia A. Earle

xi

Acknowledgments

xvii

Introduction

xix

Part 1 Geologic History of the Gulf of Mexico 1

Tectonic Evolution of the Gulf of Mexico Basin Dale E. Bird, Kevin Burke, Stuart A. Hall, and John F. Casey

2

Geology of the Florida Platform—Pre-Mesozoic to Recent Thomas M. Scott

17

3

Pre-Holocene Geological Evolution of the Northern Gulf of Mexico Basin William E. Galloway

33

4

Northern Gulf of Mexico Sea-Level History for the Past 20,000 Years James H. Balsillie and Joseph F. Donoghue

53

3

Part 2 Eastern Gulf of Mexico 5

Florida Gulf Coast Estuaries: Tampa Bay and Charlotte Harbor Gregg R. Brooks

73

6

Beaches, Barrier Islands, and Inlets of the Florida Gulf Coast Richard A. Davis

89

7

Florida Gulf of Mexico Continental Shelf—Great Contrasts and Significant Transitions Albert C. Hine and Stanley D. Locker

101

vii

viii ~ Contents

8

West Florida Continental Slope Gregg R. Brooks and Charles W. Holmes

129

9

A Review of Recent Depositional Processes on the Mississippi Fan, Eastern Gulf of Mexico David C. Twichell

141

Part 3 Northern Gulf of Mexico 10

11

12

13

Recent Geologic Framework and Geomorphology of the Mississippi-Alabama Shelf, Northern Gulf of Mexico James G. Flocks, Nicholas F. Ferina, and Jack L. Kindinger

157

Mississippi River Delta Plain, Louisiana Coast, and Inner Shelf: Holocene Geologic Framework, Processes, and Resources S. Jeffress Williams, Mark Kulp, Shea Penland, Jack L. Kindinger, and James G. Flocks

175

Development of the Northwestern Gulf of Mexico Continental Shelf and Coastal Zone as a Result of the Late Pleistocene-Holocene Sea-Level Rise Charles W. Holmes

195

Surficial Geology of the Northern Gulf of Mexico Continental Slope: Impacts of Fluid and Gas Expulsion Harry H. Roberts

209

14

Energy Resources of the Northern Gulf of Mexico Basin A. Curtis Huffman Jr., Peter D. Warwick, and Warren I. Finch

229

15

Gas Hydrates in the Gulf of Mexico Deborah R. Hutchinson, Carolyn D. Ruppel, Harry H. Roberts, Robert S. Carney, and Michael A. Smith

247

Part 4 Mexico 16

17

The Chicxulub Impact Crater and Its Influence on the Regional Hydrogeology in Northwest Yucatan, Mexico Mario Rebolledo-Vieyra, Luis E. Marin, Alberto Trejo-Garcia, and Virgil L. Sharpton Mexican Littoral of the Gulf of Mexico Arturo Carranza-Edwards

279 291

Part 5 Coral Reefs 18

Coral Reefs of the Florida Keys: Late Quaternary Stratigraphy Barbara H. Lidz

299

19

Florida Middle Ground Reef Complex Gregg R. Brooks and David Mallinson

331

20

Mexican Coral Reefs W. David Liddell and John W. Tunnell Jr.

341

21

Habitat-forming Deepwater Scleractinian Corals in the Gulf of Mexico W. W. Schroeder and S. D. Brooke

355

Contents ~ ix

Part 6 Processes 22

23

Planktic Foraminiferal Relative Abundance and Trends in Gulf of Mexico Holocene Sediments: Records of Climate Variability R. Z. Poore, S. Verardo, J. Caplan, K. Pavich, and T. Quinn Over 300 Years of Anthropogenic and Naturally Induced Low- Oxygen Bottom-Water Events on the Louisiana Continental Shelf Lisa E. Osterman, Richard Z. Poore, Peter W. Swarzenski, David Hollander, and R. Eugene Turner

367

381

Appendix 1

391

Appendix 2

401

Appendix 3

411

List of Contributors

433

Index

439

Foreword: Fifty-Year Update of Bulletin 89

Just over 50 years ago, a group of prominent marine scientists of their day agreed to begin work on a digest of existing knowledge on the Gulf of Mexico. The effort was proposed by Lionel A. Walford of the U.S. Fish and Wildlife Service and Waldo L. Schmitt of the U.S. National Museum of Natural History, during a meeting of the Gulf and Caribbean Fisheries Institute in Miami. Paul S. Galtsoff of the Fish and Wildlife Service agreed to coordinate the project, the magnitude of which he subsequently found far exceeded his expectations. However, 3 years of effort by 55 contributors and additional months of editing resulted in the 1954 publication of a classic reference work entitled Gulf of Mexico—Its Origin, Waters, and Marine Life as Fishery Bulletin 89, Fishery Bulletin of the Fish and Wildlife Service, volume 55 (Galtsoff 1954). The table of contents for the volume appears at the end of this foreword. On the title page of the work is an explanatory note that it was “Prepared by American scientists under the sponsorship of the Fish and Wildlife Service, United States Department of the Interior” and that the effort was “Coordinated by Paul S. Galtsoff,” who is generally indicated as the editor in bibliographic references to it. For more than 50 years this reference volume—commonly referred to simply as “Bulletin 89” by hosts of marine scientists, agency personnel, and students familiar with it—has provided a benchmark on which to build. Chapters on the history of exploration, geology, meteorology, physical and chemical oceanography, biota, and pollution remain extremely valuable as reference works,

some now primarily for historical context. Counted among the contributors were the most distinguished North American marine scientists of their day, and visibility for a number of them was further enhanced by the extensively cited chapters they contributed to this volume. The group included the most qualified federal agency scientists, museum curators, marine laboratory investigators, and university professors who could be assembled. It broadly represented taxonomic authorities selected to cover almost every possible biotic group, with acknowledged omission of some groups for which willing expertise could not be found. The original Bulletin 89 was heavily slanted toward biology, reflecting the focus of that era. A page count by topic reveals 63% biology (plant and animal communities, 10%; biota, 53%), oceanography 11%, geology 9%, history 6%, pollution 4%, meteorology 2%, and the index 5% (see table of contents). At the time of this writing, only one of the 55 original contributors remains alive. However, all the original contributors, and especially the far larger number of students they mentored, have contributed to a massive body of information on the Gulf of Mexico since 1954. In addition to this core group, a number of other workers—many now in laboratories, agencies, and university programs that did not exist 50 years ago—have made tremendous contributions to the baseline knowledge of the Gulf of Mexico since publication of the original volume. In September 2000, Ed Harte, former owner of the

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Corpus Christi Caller-Times and Harte-Hanks Publishing, gave Texas A&M University–Corpus Christi (TAMU– CC) a $46 million endowment to establish a research institute to study and conserve the Gulf of Mexico. Soon afterward, then President Robert Furgason obtained an additional $18 million from the State of Texas for a building to house the institute. Sylvia Earle, whose book Sea Change (Earle 1995) had inspired Harte’s gift, was invited to chair the Advisory Council. She and Bob Furgason then began establishing a world-class Advisory Council of leaders in science, academics, conservation, government, and industry. John W. (“Wes”) Tunnell Jr. was subsequently asked to serve as associate director to assist in guiding the institute development process, to coordinate construction of the new building, and to develop a new doctoral-level graduate program with other TAMU–CC faculty. This newly developing organization was given the name Harte Research Institute (HRI) for Gulf of Mexico Studies (Tunnell and Earle 2004). Further information about HRI and TAMU–CC can be found at their respective websites (www.harteresearchinstitute.org and www .tamucc.edu). After 2 HRI Advisory Council meetings, Wes Tunnell was encouraged to develop some “early” projects during the formative years of HRI in order to get a jump start on its mission of developing a cooperative and collaborative research institute focused on the long-term sustainable use and conservation of the Gulf of Mexico. From that emerged a multi-year, tri-national initiative called The Gulf of Mexico—Past, Present, and Future (Tunnell et al. 2004). The initiative had 9 components, all of which included participation by the 3 countries surrounding the Gulf of Mexico: Cuba, Mexico, and the United States. Three of the 9 components centered on determining current knowledge about the Gulf of Mexico ecosystem: (1) State of Knowledge Workshop, (2) Biodiversity of the Gulf of Mexico Project (Tunnell 2005), and (3) preparation of a 50-year update of Bulletin 89. The biodiversity and Bulletin 89 projects were initially conceived and discussed by Sylvia Earle, Wes Tunnell, and Darryl Felder in late 2001 and early 2002. Concept development continued through early 2003, when a steering committee was formed to develop ideas further and establish an implementation strategy. Steering Committee members included Fernando Alvarez, Bill Bryant, Ernesto Chávez, Luis Cifuentes, Steve Dimarco, Quenton Dokken, Sylvia Earle (co-chair), Elva Escobar, Ernie Estevez, Darryl Felder, Suzanne Fredericq, María Elena Ibarra, Chuck

Kennicutt, Paul Montagna, Marion Nipper, Worth Nowlin, Manuel Ortiz, David Pawson, Nancy Rabalais, Wes Tunnell (chair), and Gene Turner. Overall objectives for the projects were as follows: • to produce an updated Bulletin 89 • to provide a benchmark work by the leaders in the field at the beginning of the 21st century • to provide a synthesis of all work to date to the scientific, management, business, and policy communities to encourage an ecosystem view of the Gulf of Mexico • to encourage cooperation and collaboration among U.S., Mexican, and Cuban scientists working in the Gulf of Mexico, and • to determine information gaps in knowledge of the Gulf of Mexico, so targeted research can be encouraged to fill those gaps. The State of Knowledge Workshop, held 14–16 October 2003, became the kickoff for the biodiversity and Bulletin 89 projects. The 50-year update of Bulletin 89 grew from one volume to 7, broadly including geology, physical and chemical oceanography, biota, anthropogenic issues, ecosystem-based management, and socioeconomics. Likewise, the effort was expanded from 55 authors in 1954 to more than 200 for the new effort. Most of the knowledge gained and presented in the original Bulletin 89 was from research cruises and expeditions to the Gulf during the late 19th and early 20th centuries and from a few fledgling marine science labs and oceanography programs started in the early to mid-1900s, but massive efforts have followed in their wake. At the dawning of a new century, researchers at marine labs and universities encircle the Gulf in Cuba, Mexico, and the United States (see www.gulfbase.org), and instrumentation, technology, and communication have greatly expanded our knowledge of the Gulf. The U.S. Environmental Protection Agency’s Gulf of Mexico Program has identified priority problems affecting the northern Gulf of Mexico, and the agency recently published the research needs of that region (EPA 2002). A network of United Nations organizations has declared the Gulf of Mexico as one of 64 large marine ecosystems in the world (Kumpf et al. 1999). The U.S. Commission on Ocean Policy released its report, An Ocean Blueprint for the 21st Century, in August 2004, listing 212 recommendations for actions to care for and manage U.S. coasts and oceans better. President George W. Bush subsequently

Foreword ~ xiii

issued his Ocean Action Plan response in December 2004, singling out the Gulf of Mexico as a region of special concern. Next, HRI sponsored the first State of the Gulf of Mexico Summit. This 3-day summit, held 26–30 March 2006, was attended by 450 invited guests from the United States and Mexico, and it focused on governance, catastrophic events, sustainability, economics, public health, and the environment (Tunnell and Dokken 2006). Future summits are planned at regular intervals, perhaps every 3 years, at key locations around the Gulf of Mexico. While the challenge was daunting, an update of Bulletin 89 was long overdue. As the 50th anniversary of its publication has passed, the range and scope of primary literature sources on the Gulf of Mexico have become so expansive as to be all but unmanageable for most workers. For almost all subject areas, no authoritative digests centered on the Gulf of Mexico have appeared since Bulletin 89. Yet many treatments in that work are clearly outdated and are of limited value except as historical starting points. Furthermore, it was urgent to begin compilations for this updated digest before the marine science community sustained further loss of continuity in expertise. We have already lost all but one of the original contributors to Bulletin 89, the passage of 50 years has claimed a large number of the subsequent generation of workers, and others are late in their careers. This is perhaps most evident in what has become a very limited pool of qualified systematists to draw upon for expertise concerning diversity and taxonomy of the Gulf of Mexico biota. It is noteworthy that one remarkable scientist who contributed to the original effort, the late Frederick M. (Ted) Bayer, also coauthored a chapter for one volume (Biota) immediately prior to his passing. We collectively thank all contributors to this immense effort, especially the editors and coordinators of each volume for their multi-year commitments to this project. Many users will benefit for decades to come because of their dedicated efforts. We especially thank Ed Harte for putting in place the infrastructure needed to undertake a project of this magnitude, and we offer these volumes as an early step in response to his charge—“make a difference.” John W. Tunnell Jr. Darryl L. Felder Sylvia A. Earle Bulletin 89: Fifty-Year Update Series Coordinators

Bulletin 89 Table of Contents (Abbreviated) Gulf of Mexico—Its Origin, Waters, and Marine Life Paul S. Galtsoff, Coordinator Fishery Bulletin 89, 1954 I. Historical sketch by Paul S. Galtsoff (34 pp.) II. Geology Shorelines and coasts of the Gulf of Mexico by W. Armstrong Price (27 pp.) Geology of the Gulf of Mexico by S. A. Lynch (20 pp.) III. Marine meteorology by Dale F. Leipper (10 pp.) IV. Physics and Chemistry Tides and sea level in the Gulf of Mexico by H. A. Marmer (18 pp.) Physical oceanography of the Gulf of Mexico by Dale F. Leipper (19 pp.) Light penetration in the Gulf of Mexico by William S. Shoemaker (3 pp.) Distribution of chemical constituents by Robert H. Williams (9 pp.) The recovery of minerals by C. M. Shigley (7 pp.) V. Plant and animal communities Phytoplankton by Charles C. Davis (7 pp.) Zooplankton by Hilary B. Moore (2 pp.) Red tide by Reuben Lasker and F. G. Walton Smith (4 pp.) Marine algal vegetation of the shores by William Randolph Taylor (16 pp.) Flowering plants by Robert F. Thorne (10 pp.) Bottom communities by Joel W. Hedgpeth (12 pp.) VI. Bacteria, fungi, and unicellular algae Marine bacteria and fungi by Claude E. ZoBell (6 pp.) Dinoflagellates by Herbert W. Graham (4 pp.) Diatoms by Paul Conger (6 pp.) VII. Protozoa Foraminifera by Fred B. Phleger and Frances L. Parker (7 pp.) Protozoa by Victor Sprague (14 pp.) VIII. Sponges, coelenterates, and ctenophores Porifera by J. W. Tierney (3 pp.) Commercial sponges by F. G. Walton Smith (4 pp.) Hydroids by Edward S. Deevey Jr. (6 pp.) Hydromedusae by Mary Sears (2 pp.) Siphonophores by Mary Sears (2 pp.) Scyphozoa by Joel W. Hedgpeth (2 pp.) Anthozoa: Alcyonaria by Frederick M. Bayer (6 pp.) Anthozoa: The anemones by Joel W. Hedgpeth (6 pp.) Madreporaria by F. G. Walton Smith (5 pp.) Ctenophores by Mary Sears (1 p.)

xiv ~ Foreword

IX. Free-living flatworms, nemerteans, nematodes, tardigrades, and chaetognaths Free-living flatworms (Turbellaria) by L. H. Hyman (2 pp.) Nemerteans by Wesley R. Coe (7 pp.) Echinoderida by B. G. Chitwood (1 p.) Free-living nematodes by B. G. Chitwood and R. W. Timm (11 pp.) Tardigrades by B. G. Chitwood (1 p.) Chaetognatha by E. Lowe Pierce (3 pp.) X. Parasitic worms Parasitic helminths by Asa C. Chandler and Harold W. Manter (2 pp.) Trematoda by Harold W. Manter (16 pp.) Cestoda by Asa C. Chandler (3 pp.) Acanthocephala by Asa C. Chandler (1 p.) Nematoda by Asa C. Chandler (2 pp.) XI. Bryozoa, Brachiopoda, Phoronida, and Enteropneusta The Bryozoa by Raymond C. Osburn (2 pp.) Brachiopoda by G. Arthur Cooper (3 pp.) Phoronida by Joel W. Hedgpeth (1 p.) Enteropneusta by Joel W. Hedgpeth (1 p.) XII. Echinoderms Echinoderms (other than holothurians) by Austin H. Clark (7 pp.) Holothurians by Elisabeth Deichmann (30 pp.) XIII. Annelids and miscellaneous worms Polychaetous annelids by Olga Hartman (5 pp.) Miscellaneous vermes by Joel W. Hedgpeth Echiurid (1 p.) Sipunculida (1 p.) XIV. Arthropods Xiphosura by Joel W. Hedgpeth (1 p.) Pycnogonida by Joel W. Hedgpeth (3 pp.) Ostracoda by Willis L. Tressler (9 pp.) Copepoda by Waldo L. Schmitt (4 pp.) Cirripedia (barnacles) by Doral P. Henry (4 pp.) Mysidacea and Euphausiacea by Albert H. Banner (2 pp.) Stomatopoda by Fenner A. Chace, Jr. (2 pp.) Decapoda by Ellinor H. Behre (5 pp.) Commercial shrimps by Milton J. Lindner and William W. Anderson (5 pp.) Spiny lobster by F. G. Walton Smith (3 pp.) XV. Mollusks Mollusks by Harald A. Rehder (6 pp.) Cephalopoda by Gilbert L. Voss (4 pp.) Oyster Biology by Philip A. Butler (11 pp.)

Oyster reefs by W. Armstrong Price (1 p.) XVI. Tunicates and lancelets Tunicata by Willard G. Van Name (3 pp.) Lancelets by Joel W. Hedgpeth (1 p.) XVII. Fishes and sea turtles Fishes by Luis R. Rivas (3 pp.) Commercial fishes by George A. Rounsefell (6 pp.) Sea turtles by F. G. Walton Smith (3 pp.) XVIII. Birds George H. Lowery, Jr., and Robert J. Newman (22 pp.) XIX. Mammals Gordon Gunter (9 pp.) XX. Pollution of water (21 pp.)

References Earle, S. 1995. Sea Change: a Message of the Oceans. New York: Ballantine Books. 384 pp. EPA (Environmental Protection Agency). 2002. Critical Scientific Research Needs Assessment for the Gulf of Mexico. Prepared by the Research Subcommittee of the Monitoring, Modeling, and Research Committee for the Gulf of Mexico Program Office. 47 pp. Galtsoff, Paul S. 1954. Gulf of Mexico—Its Origin, Waters, and Marine Life. Fishery Bulletin 89, U.S. Fish and Wildlife Service, vol. 55. Washington, D.C. 604 pp. Kumph, H., K. Steidinger, and K. Sherman (eds.). 1999. The Gulf of Mexico Large Marine Ecosystem. Malden, Mass.: Blackwell Science. 626 pp. Tunnell, J. W., Jr. 2005. Biodiversity of the Gulf of Mexico project. Pp. 285–86 in P. Miloslavich and E. Klein (eds.), Caribbean Marine Biodiversity: The Known and the Unknown. Lancaster, Penn.: DEStech Publications. Tunnell, J. W., Jr., and Q. R. Dokken (eds.). 2006. Proceedings of the State of the Gulf of Mexico Summit. Corpus Christi, Texas, 28–30 March 2006. Corpus Christi, Tex.: Harte Research Institute for Gulf of Mexico Studies, Texas A&M University–Corpus Christi. 44 pp. Tunnell, J. W., Jr., and S. A. Earle. 2004. Harte Research Institute for Gulf of Mexico Studies: Initiatives in Marine Science Research. Pp. 132–41 in R. L. Creswell (ed.), Proceedings of 55th Annual Gulf and Caribbean Fisheries Institute, Xel-Ha, Quintana Roo, Mexico. Fort Pierce, Fla.: Gulf and Caribbean Fisheries Institute. Tunnell, J. W., Jr., D. L. Felder, and S. A. Earle. 2004. El Golfo de México-pasado, presente, y futuro: una colaboración entre Estados Unidos de América, México y Cuba. Pp. 361–71 in M. Caso-Chávez, I. Pisanty, and E. Ezcurra (eds.), Diag-

Foreword ~ xv

nóstico ambiental del Golfo de México. Instituto Nacional de Ecología, (INECOL A.C.) and Harte Research Institute for Gulf of Mexico Studies TAMU–CC, 2 vols. México City: Secretaría de Medio Ambiente y Recursos Naturales (SEMARNAT). U.S. Commission on Ocean Policy. 2004. An Ocean Blueprint

for the 21st Century. Final Report. Washington, D.C.: U.S. Commission on Ocean Policy. 676 pp. U.S. Ocean Action Plan: The Bush Administration’s Response to the U.S. Commission on Ocean Policy. 2004. Washington, D.C. 39 pp.

Acknowledgments

Contemplation of this geology volume was initiated in 2005 and authors began submission of their manuscripts in 2006. It has taken four years of tremendous effort by many people to reach publication. First and foremost, the contents of this book would not have been complete or possible were it not for the contributions of the many geologists whose significant research and years of hard work both in the field and office is summarized on the following pages. We gratefully acknowledge each of the authors and coauthors who took time from their institutional duties or retirement. Their efforts and patience will forever enhance our understanding of the formation, depositional history and processes, corals, and climate history of the Gulf of Mexico. We sincerely thank Wes Tunnell for his assistance and insight to initiate this Gulf of Mexico volume and the other similar volumes in this series. He has been wonderful to work with throughout the process and has provided

positive influence and direction toward completion of the volume. In addition, we would also like to acknowledge the years of support for this book by the U.S. Geological Survey and, in particular, John Haines, Lisa Robbins, Dawn Lavoie, and Jack Kindinger. Because of the need to use color to best represent the data in the illustrations we gratefully acknowledge contributions from the U.S. Geological Survey. Included in our list of thanks must be Gene Shinn, who graciously reviewed this volume and gave wonderful advice. We also thank the editing staff of the Harte Research Institute, especially Susan Shirley for her exceptional efforts on final edits of all manuscripts. Also, Shannon Davies from Texas A&M University Press, who assisted us in the intricacies of publication and Chandra Dreher (USGS) for help with final proofs. And finally, we warmly recognize and thank Barbara Lidz, whose input, efforts, and words of advice were invaluable.

xvii

Introduction

“No rock, no water, no ecosystem” is a phrase used repeatedly by Eugene Shinn at the University of South Florida College of Marine Science. The phrase highlights the fact that to understand a region as vast as the Gulf of Mexico, one must know the environmental foundation of the area. According to data stored in the American Geological Institute database GEOREF, there have been over 10,000 journal publications on various aspects of the geology of the Gulf of Mexico. Whereas about 25% of these have been published by industry, almost 70% have been produced by 4 major public entities: the University of Texas, Texas A&M University, Louisiana State University, and the U.S. Geological Survey. The remaining 5% are products of various public entities throughout the world. Although this volume of data is significant, it is but a small fraction of that residing in the files of various petroleum and minerals organizations. Some of these data are public and are available for future research; however, most remain sequestered in the files of a highly competitive industry.

Geologic Research in the Gulf of Mexico The modern era of geologic research started in the beginning of the second half of the 20th century with the initiation of American Petroleum Institute (API) Project 51 and the formation of the Department of Oceanography at Texas A&M University. With initial professional appointments in physical, biological, chemical, geological, and meteorological oceanography, the Texas A&M

University program recognized from the beginning the inherent multidisciplinary nature of the marine science field. Operations began in September 1949 with a curriculum leading to graduate degrees in oceanography. The first geological oceanographer at Texas A&M University was W. Armstrong Price, and the first PhD graduate from the department was Warren C. Thompson (1953), whose dissertation was titled “Geological Oceanography of Atchafalaya Bay.” At the same time, other institutions became interested in the Gulf of Mexico. The first regionally comprehensive studies began under the guidance of Francis P. Shepard of Scripps Institute of Oceanography. His investigation of sediment distribution came about as a result of many years of deliberation by committees of the American Petroleum Institute and the American Association of Petroleum Geologists, all of which led to creation of API Project 51. The objective of Project 51 was to study modern sediment on the northwestern margin of the Gulf of Mexico. In addition to describing the distribution and facies relations of various types of sediment, the project established procedures that have become the standard for all subsequent studies. Details of these procedures are described in the preface to the summary volume Recent Sediments, Northwest Gulf of Mexico (Shepard et al. 1960). Early in the 1950s, other discoveries led to more detailed investigations. For example, during a cruise in 1954 across the Gulf of Mexico, subaqueous hills were discovered on what was then thought to be the featureless Sigsbee Plain (Ewing et al. 1958). The hills, thought xix

xx ~ Introduction

to be salt domes, were verified in 1968 at Site 2 of the first leg of the Deep-Sea Drilling Program (DSDP) of the R / V Glomar Challenger. Between 1950 and 1970, many other studies were carried out to resolve specific problems. Although the studies were excellent, none were gulfwide in scale. In 1969, the USGS and U.S. Naval Oceanographic office initiated a cooperative geophysical-geological program. Data collection began in September 1969, yielding more than 25,000  km of continuous seismic-reflection profiles, which provided a clean picture of the complex structure of the Gulf of Mexico (Garrison and Martin 1973). For part of this program, experimental satellite navigation was used in conjunction with Loran-C. The geologic phase of this program produced many sediment cores throughout the Gulf. Sediment from the cores formed the basis of many studies of the upper Pleistocene and Holocene (including Bouma 1972; Davies 1972; Kennett and Huddleston 1972; Holmes 1976). Later in the 1980s the USGS Geological Long-Range Incline Asdic (GLORIA) side-scan survey and the Minerals Management Service (MMS) outer continental shelf (OCS) environmental studies produced large amounts of high-quality data that provided additional knowledge of the region. The results of all of these studies are included as chapters within this volume. Although it is believed that this volume differs from other summaries because it contains information on each region of the Gulf of Mexico, there are areas for which information remains scarce, including the south and central portions of the basin, the Sigsbee Deep, the Gulf of Campeche, and the Yucatan Channel. In 1982, the United Nations established the Exclusive Economic Zone (EEZ), which divided the Gulf of Mexico in half. The southern portion came under the exclusive jurisdiction of the Republic of Mexico. As a result, much of the data in the south that have been collected since then have remained propriety property of the Mexican government. However, before establishment of the EEZ, some regional cooperative studies were conducted in the southern Gulf of Mexico. In addition to the gulfwide data collected by the USGS and U.S. Navy in 1969, an additional study was carried out under the auspices of the 1971 International Decade of Ocean Exploration program which collected acoustical-reflection profiles in the Gulf of Campeche. The research on the “Mexican Ridges” near Veracruz, Mexico, (Moore 1972) and the results of these two studies were summarized by Massingill et al. (1973). A few scientific cruises into the Sigsbee Basin

have produced intriguing results, such as the discovery of asphalt-like pavement covering large areas of the ocean floor, but the region remains relatively unexplored. Similarly, the deep water between the Yucatan Channel and the Florida Peninsula also remains a mare incognitum. Although great strides have been made in understanding Gulf of Mexico geology, there are still significant gaps in our knowledge of exactly how this very important basin, with its rich mineral and biological resources, formed. Because of these resources there is now a renewed interest in virtually all aspects of gulfwide geological investigations. Some of the most prominent projects are the mapping of the west Florida shelf and the geologic investigation of Pulley Ridge, a deepwater reef off southwest Florida. This reef exists in over 100 m of water but is populated with organisms normally found in shallow water. This reef grew on a drowned barrier island chain. As a result of the hurricanes that devastated the northern Gulf of Mexico in 2005, there is a renewed interest in defining the roles of barrier islands and their protection, or lack thereof, of the mainland of this vulnerable coast. Information on these investigations will be applicable to the entire Texas–Louisiana coastline. Further studies are underway to resolve the forces that have led to the formation of hypoxic zones throughout the region, and the geologic aspects of the formation of harmful algal blooms (HABs). We also need to understand the role of gas hydrates and how they affect sediment stability. This information will be very important as the petroleum industry moves into deeper and deeper water. Thus, while the innovation of geologic models that have been established by advances in technology has been significant, these advances have raised more questions that need to be answered.

Purpose and Contents of this Volume The number of publications on the Gulf of Mexico increased exponentially since the first “Bulletin 89” was produced. Publishing activity reached a peak during 1995–2000, but after 2000 the number of publications waned. There are many reasons for this decline, including the closing of several research laboratories, the decline in industrial employment, declines in student enrollment, and finally the retirement of many scientists engaged in Gulf of Mexico research. Nowhere is there a single publication with an integrated synthesis and interpretation encompassing the

Introduction ~ xxi

entire Gulf of Mexico. The closest is The Gulf of Mexico Basin (Salvador 1991), which touched on various aspects of basin physiography but focused mainly on the basin as a petroleum source. The objective of this present volume, while including geologic origin and Late Mesozoic and Tertiary history of the basin, is to focus on the processes across the entire basin that have been active during the Quaternary. This volume has been planned as a systematic, cohesive, and, we hope, comprehensive description of the geology of the Gulf of Mexico basin, not a collection of independent papers brought together because they address a common area. The volume is divided into 6 parts. The first part contains 4 chapters dealing with basin origin. The first chapter, by Bird and others, focuses on deep geophysical information and presents a model for the basin’s origin. The second chapter, by Scott, updates the most recent information on Upper Mesozoic–Tertiary strata of the Florida Peninsula. In the third chapter, Galloway presents a model of pre-Holocene geological evolution of the northwestern Gulf of Mexico. The last chapter in this section by Balsillie and Donoghue addresses the important and timely record of late Pleistocene–Holocene sealevel rise. The second part presents 5 chapters dedicated to the geology of the eastern Gulf of Mexico. The first chapter in this part, by Brooks, is a summary of information on estuaries of the central Florida Peninsula. The second chapter, by Davis, presents significant information on the formation and processes that led to the formation of beaches, barrier islands, and inlets along the central Florida coast. The third chapter, by Hine and Locker, is a 50-year summary of data gathered on the west Florida shelf. This very important region is just coming of age because it was all but forgotten until the 1980s. The fourth chapter, by Brooks and Holmes, completes the picture of this region, detailing the development of the west Florida slope, and lastly, the chapter by Twichell presents a detailed exposé of the development of the lower Mississippi fan in the eastern Gulf of Mexico. Part 3 of this volume focuses on important Holocene sedimentary regions of the northwestern Gulf of Mexico. The first chapter in this part, by Flocks and others, summarizes development of the Mississippi–Alabama shelf. The second chapter, by Williams and others, explores the significant information about the Mississippi Delta and its importance to the geology of the Gulf of Mexico. The Mississippi River and its evolving deltas have had a dominant role in the development of the clastic strata that exist

on both sides of its present mouth. The late Pleistocene– Holocene development of the region to the west is described by Holmes in the third chapter. The fourth chapter, by Roberts, details the processes that are active on the upper Louisiana shelf. This region is vital to petroleum exploration, and it is a region where sediment can remain unstable due to a combination of deposition and salt tectonics. The fifth chapter, by Huffman and others, is a summary of the resources in the Gulf of Mexico basin. The final chapter, by Hutchinson and others, details an extremely interesting and little-understood phenomenon in the Gulf of Mexico: gas hydrate. Understanding how gas hydrate affects sediment stability and how it forms and vanishes is explained in their chapter. The fourth part in the volume presents information from the southern Gulf. This area of the basin is little understood because much of the data is property of PEMEX (Petróleos Mexicanos, the Mexican national petroleum company). The first chapter, by RebolledoVierya and others, details geology in one of the most interesting locations on Earth, the impact crater at Chicxulub that formed at the end of the Cretaceous when a meteorite impact reportedly extinguished the dinosaurs. The second chapter, by Carranza-Edwards, presents data on beach formations along the southern Mexican coast. The fifth part of the volume focuses on reefs that have formed around the Gulf of Mexico. The best-studied fossil reef is that which formed the Florida Keys at the southern tip of Florida. In the first chapter in this part, Lidz summarizes 5 decades of geologic exploration of the Florida Keys and its modern reef ecosystem. The second chapter, by Brooks and Mallinson, details the enigmatic socalled reef on the north-central Florida shelf—the Florida Middle Grounds. Liddell and Tunnell present data on the reefs and extensive carbonate sediments that formed the Yucatan Peninsula in the third chapter. In the last chapter, Schroeder and Brooke provide geologic data on the little-known and poorly understood deep reef-like features. These topographic features are known to extend from Florida to the western edge of the Louisiana shelf and, at present, are being investigated to define their value as a biological resource. The last part of the volume details processes of climate change that have had profound effects on the present environment in the basin. In the first chapter in this part, Poore and others present data on the climate changes as recorded by the presence of planktic foraminifera species in the sediments. Their chapter relates foraminiferal variation in the sediment record to climate changes on the

xxii ~ Introduction

adjacent continent for the last century. The final chapter, by Osterman and others, focuses on hypoxia, a condition now recognized as a major biological hazard on the inner shelf in the northern Gulf. Their study reveals that hypoxic episodes have occurred previously, although not as widespread as in recent times. Chapters in this volume represent an updated, basinwide compilation of scientific knowledge pertaining to the geology of the Gulf of Mexico. With long hours of work and effort from authors and editors, we hope the contents will serve as an inspiration and a lasting resource for Gulf of Mexico researchers for many years to come.

References Bouma, A. H. 1972. Distribution of sediments and sedimentary structures in the Gulf of Mexico. Pp. 35–52 in R. Rezak and V. E. Henry (eds.), Contribution on the Geological and Geophysical Oceanography of the Gulf of Mexico. Houston, Tex.: Gulf Publishing, Inc. Davies, D. K. 1972. Mineralogy, petrology, and derivation of sands and silts of the continental slope, rise, and abyssal plain of the Gulf of Mexico. Journal of Sedimentary Petrology 42:59–65. Ewing, M., D. B. Ericson, and B. C. Heezen. 1958. Sediments and topography of the Gulf of Mexico. Pp. 995–1053 in L. G. Weeks (ed.), Habitat of Oil. Tulsa, Okla.: American Association of Petroleum Geologists.

Garrison, L. E., and R. G. Martin. 1973. Geologic Structures in the Gulf of Mexico basin. U.S. Geological Survey Professional Paper 773. Reston, Va.: U.S. Geological Survey. 85 pp. Holmes, C. W. 1976. Distribution, Regional Variation, and Geochemical Coherence of Selected Elements in the Sediment of the Central Gulf of Mexico. U.S. Geological Survey Professional Paper 928. Reston, Va.: U.S. Geological Survey. 26 pp. Kennett, J. P., and P. Huddleston. 1972. Late Pleistocene paleoclimatology, formaminiferal biostratigraphy, and tephrachronology, western Gulf of Mexico. Quaternary Research 2:45–62. Massingill, J. V., R. N. Bergantino, H. S. Fleming, and R. H. Feden. 1973. Geology and genesis of the Mexican ridges: western Gulf of Mexico. Journal of Geophysical Research 78:225–38. Moore, G. W. 1972. Acoustic-Reflection Profiles: Bay of Campeche. International Decade of Ocean Exploration, USGS-GD-72–002.Washington, D.C.: U.S. Geological Survey. 21 pp. Salvador, A. (ed.). 1991. The Geology of North America, Volume J, Gulf of Mexico Basin. Boulder, Colo.: Geological Society of America. 568 pp. Shepard, F. P., F. B. Phleger, and T. H. Van Andel (eds.). 1960. Recent Sediments, Northwest Gulf of Mexico. Tulsa, Okla.: American Association of Petroleum Geologists. 394 pp. Thompson, W. C. 1953. Geological oceanography of Atchafalaya Bay. Ph.D. dissertation, Texas A&M University, College Station. 152 pp.

Gulf of Mexico Origin, Waters, and Biota

Part 1

␥␥␥

Geologic History of the Gulf of Mexico

␥1

Tectonic Evolution of the Gulf of Mexico Basin Dale E. Bird, Kevin Burke, Stuart A. Hall, and John F. Casey

The formation of the Gulf of Mexico basin was preceded by the Late Triassic breakup of Pangea, which began with the collapse of the Appalachian Mountains (ca. 230 Ma; Dewey 1988). Gondwanan terranes of the southern part of the Gulf States, eastern Mexico, and the Yucatan Peninsula remained sutured onto the North American continent as it drifted away from the African-Arabian-South American continent (or Residual Gondwana, Burke et al. 2003). Early seafloor spreading in the central Atlantic Ocean, from about 180 Ma to 160 Ma, included 2 jumps of the seafloor-spreading center to new locations. The timing of the latter ridge jump (ca. 160 Ma) correlates with initial rifting and rotation of the Yucatan block. The Gulf of Mexico ocean basin is almost completely bounded by continental crust. Its shape requires that at least one ocean-continent transform boundary was active while the basin was opening (Fig. 1.1). Evolutionary models differ between those that require the basin to open by rotation along a single ocean-continent transform boundary (counterclockwise rotation of the Yucatan block), and those that require the basin to open by rotation along a pair of subparallel ocean-continent transform boundaries (essentially northwest-southeast motion of the Yucatan block). Although many models have been proposed, most workers now agree that counterclockwise rotation of the Yucatan Peninsula block away from the North American Plate, involving a single ocean-continent transform boundary, led to the formation of the basin; many of these workers suggest that this rotation occurred between 160 Ma (Oxfordian) and 140 Ma (Berriasian-Valanginian)

about a pole located within 5° of Miami, Florida (Humphris 1979; Shepherd 1983; Pindell 1985, 1994; Dunbar and Sawyer 1987; Salvador 1987, 1991; Burke 1988; Ross and Scotese 1988; Christenson 1990; Buffler and Thomas 1994; Hall and Najmuddin 1994; Marton and Buffler 1994). Evidence cited for this model of basin evolution includes: (1) paleomagnetic data from the Chiapas massif of the Yucatan Peninsula (Gose et al. 1982; Molina-Garza et al. 1992), (2) fracture zone trends interpreted from magnetic data (Sheperd 1983; Hall and Najmuddin 1994), (3) non-rigid tectonic reconstructions (Dunbar and Sawyer 1987; Marton and Buffler 1994), and (4) kinematic reconstructions making use of geological constraints, well data, and geophysical data such as seismic refraction, gravity, and magnetics (Pindell 1985, 1994; Christenson 1990; Marton and Buffler 1994). Most workers consider the total counterclockwise rotation of the Yucatan block to be between 42° and 60° (Dunbar and Sawyer 1987; Ross and Scotese 1988; Hall and Najmuddin 1994; Marton and Buffler 1994; Schouten and Klitgord 1994). Differences in the amount of rotation reflect the close proximity of the Yucatan block to its pole of rotation. That is, a small change in this distance can produce a relatively large change in the rotation angle when the plate being rotated is close to, or contains, the rotation pole. Additional support for counterclockwise rotation is provided by paleomagnetic data (Gose et al. 1982; Molina-Garza et al. 1992). The amount of counterclockwise rotation reported by these authors, 75° (Molina-Garza et al. 1992) and 130° (Gose et al. 1982), is 3

4 ~ Bird, Burke, Hall, and Casey

Figure 1.1. Gulf of Mexico basin. The bathymetry contour interval is 500 m. Keathley Canyon (KC) and Yucatan Parallel (YP) free air gravity anomaly outlines show locations of interpreted hotspot tracks. The Tamaulipas-Golden Lane-Chiapas (TGLC) free air gravity anomaly is interpreted to be produced by a marginal ridge. The extent of present-day salt deposits is shaded green (after Martin 1980). OCB is the ocean-continent boundary.

with respect to the magnetic north pole and represents a somewhat larger but more poorly determined rotation of the Yucatan block. Since 42° is roughly twice the rotation that we interpret for seafloor spreading, and this amount brings the Yucatan into a reasonable position after reconstruction, we use this estimate (Marton and Buffler 1994) for our reconstruction. Prominent basement features within the Gulf of Mexico basin are interpreted to be hotspot tracks that were created by a single mantle plume as the basin opened (Bird et al. 2005a). During the seafloor-spreading phase, this Late Jurassic mantle plume (ca. 150 Ma to 140 Ma) may have generated the hotspot tracks on the North American Plate and the Yucatan block. The tracks are identified from deep-basement structural highs that have

been mapped by integrating seismic refraction and gravity data. They are associated with high-amplitude, distinct gravity anomalies that provide the basis for a plate tectonic reconstruction that restores the western ends of the hotspot tracks with a 20° clockwise rotation of the Yucatan block, or almost one-half the total rotation required to open the Gulf of Mexico basin (Figs. 1.1, 1.2). The duration of track generation is estimated to have been about 10 Myr, or almost one-half the total time required to open the Gulf of Mexico basin. One gravity anomaly is centered over the Keathley Canyon concession area and is here called the Keathley Canyon anomaly. The second anomaly, which curves for about 630 km concentric with the Yucatan Peninsula continental margin, is here called the Yucatan Parallel anomaly. A third anomaly, oriented

Tectonic Evolution ~ 5

Figure 1.2. Gulf of Mexico gravity anomalies, free air offshore, and Bouguer onshore. Hotspot tracklines over the Keathley Canyon and Yucatan Parallel anomalies (see Fig. 1.1) were based on the rotation pole (HN) described by Hall and Najmuddin (1994). The total hotspot track growth and Yucatan block rotation, during seafloor spreading, are calculated to be 10 My and 20° (italics). The seafloor-spreading center (double white lines) is schematic, and OCB is the ocean-continent boundary.

roughly north-south and concentric with the east coast of central Mexico, extends from the Rio Grande delta in the north to just offshore Veracruz in the south (Figs. 1.1, 1.2). It is related to the Tamaulipas-Golden Lane-Chiapas fracture zone defined by Pindell (1985, 1994), and it is referred to here as the Tamaulipas-Golden Lane-Chiapas anomaly. The Tamaulipas-Golden Lane-Chiapas anomaly was produced by a marginal ridge that was created along this ocean-continent transform boundary as the basin opened. The eastern flank of the marginal ridge and the northernmost, easternmost, and southernmost terminations of the hotspot tracks are interpreted to coincide with the oceanic-continental crustal boundary in the basin (Figs. 1.1, 1.2). Prior to rotation by seafloor spreading, extension of continental crust over an 8 Myr to 10 Myr interval was the result of approximately 22° of counterclockwise rotation and lithospheric thinning. Autochtho-

nous salt appears to be confined to the continental flanks of the hotspot tracks confirming that salt was deposited during continental extension and not after ocean floor had begun to form (Fig. 1.1).

Pangea Breakup From Ladinian (Middle Triassic) to Oxfordian (early Late Jurassic), early extension associated with the breakup of Pangea occurred along the Appalachian-collapse rift system (initiated ca. 230 Ma), which extends from east Greenland and the British Isles in the north, through the Appalachian Mountains of North America, to the Takatu Rift of Guyana and Brazil in South America (Burke et al. 2003). North America-Gondwana rifting continued until about 180 Ma when seafloor spreading in the central Atlantic began (Withjack et  al. 1988). During this

6 ~ Bird, Burke, Hall, and Casey

time, the short-lived (about 2 Myr) Central Atlantic Magmatic Province (CAMP) mantle plume erupted (201 Ma), producing about 60 thousand cubic kilometers of flood basalts and associated intrusions over 2.5 million square kilometers in North and South America, Africa, and even Europe (Marzoli et al. 1999). The growth of ocean basins as continents drift apart is reflected in magnetic data. Bands of linear anomalies flanking spreading centers represent episodic reverses in the polarity of the earth’s geomagnetic field. The time intervals between polarity reversals are called chrons, and they have been identified in the world’s ocean basins for the Cenozoic Era and Late Cretaceous Period (C-series: C1 to C34), and in the earlier Mesozoic Era to about 167 Ma (M-series: M0 to M41) (Gradstein et al. 2004). Because extensional rifting in passive margins essentially stops once new oceanic lithosphere is created, closing ocean basins along geomagnetic isochrons is an objective method for analyzing reconstructed continental margins.

Mesozoic chrons from M0 to M40, including several in the Jurassic Magnetic Quiet Zone (JMQZ, from ca. 167 Ma to 155 Ma, or M41 to M26), have been identified and mapped between the Atlantis and Fifteen-Twenty fracture zones on the North American Plate, and between the Atlantis and Kane fracture zones on the African Plate (Fig. 1.3A) (Bird 2004). Chron M40 (167.5 Ma) is mapped about 65 km outboard of the African S1 magnetic anomaly and its conjugate, the Blake Spur Magnetic Anomaly (BSMA), over the eastern and western flanks of the central Atlantic (Figs. 1.3B, 1.3C). Another pair of conjugate anomalies, the S3 magnetic anomaly and East Coast Magnetic Anomaly (ECMA), are respectively located about 30  km and 180  km inboard of the S1-BSMA pair. For that reason the shift in the seafloor-spreading center, or “ridge jump,” about 90 km to the east between the BSMA and the ECMA anomalies at about 170 Ma ( Vogt et al. 1971) is supported by this study. Between the Atlantis and Kane fracture zones the width of the African JMQZ

A)

Figure 1.3. (A) Central Atlantic Ocean magnetic isochrons and fracture zones. The Mid-Atlantic Ridge (MAR) and main fracture zones are red; Atlantis, Kane, and Fifteen-Twenty (15–20) are fracture zones used to reconstruct the basin (Bird et al. 2005b). Bands of identified isochrons include the Cenozoic C-series that flank the MAR, then the older Cretaceous Magnetic Quiet Zone (CMQZ, no magnetic polarity reversals occurred during this time), then the Mesozoic M-Series (Muller et al. 1997). (B) and (C) Chron M40 is mapped about 65 km outboard of the conjugate Blake Spur Magnetic Anomaly (BSMA)–S1 Anomaly (Bird 2004) indicating that a ridge jump occurred between the conjugate East Coast Magnetic Anomaly (ECMA)–S3 Anomaly (ca. 170 Ma). Repeated chron M38 over the African flank, and absent over the North American flank, indicates a ridge jump. The Jurassic Magnetic Quiet Zone (JMQZ) is characterized by a relatively weak magnetic field.

B)

C)

Figure 1.3. (continued )

8 ~ Bird, Burke, Hall, and Casey

is about 70 km greater (22%) than the North American JMQZ. Inspection of magnetic anomalies over this range reveals that additional correlatable anomalies exist over Africa (Bird 2004), suggesting a second ridge jump of about 35 km to the west. Modeling results indicate that this jump occurred between 159 Ma and 164 Ma (chrons M32 and M38). These ridge jumps could have coincided with North American-Gondwana plate reorganizations including rifting of the Yucatan block away from North America and seafloor spreading in the Gulf of Mexico. The second ridge jump in the central Atlantic (ca. 160 Ma) roughly coincides with the initiation of Yucatan block rotation followed by the formation of the Gulf of Mexico (Dunbar and Sawyer 1987; Burke 1988; Ross and Scotese 1988; Salvador 1991; Buffler and Thomas 1994; Hall and Najmuddin 1994; Marton and Buffler 1994; Pindell 1994). The 2 ridge jumps described here are consistent in dimensions and duration with other ridge jumps observed around the world (Bird 2004). Ridge jumps have been documented along the Mid-Atlantic Ridge near the Ascension fracture zone (Brozena 1986), at 7 locations west of the East Pacific Rise including 2 currently underway (Luhr et al. 1986; Mammerickx and Sandwell 1986; Morton and Ballard 1986; Mammerickx et al. 1988), south of the Chilean Ridge (Mammerickx et al. 1988), and at 3 locations in the north Pacific (Mammerickx et al. 1988). Our closest North American-Gondwana fit (Fig. 1.4A) illustrates that final closure (to form pre-breakup Pangea) requires that: (1) the Yucatan block rotated over 40° clockwise from its present position to close the Gulf of Mexico, (2) the southern edge of the Florida shelf was contiguous with the Demerara Rise of South America and the Guinea Nose of Africa as suggested by Pindell and Dewey (1982), and (3) the Bahama Island chain must have formed while the central Atlantic was opening supporting the idea that the islands overlie a hotspot track, as was first suggested by Dietz (1973). That track is recognized here to be that of the Early Jurassic Central Atlantic Magmatic Province (CAMP) mantle plume that initially erupted at 201 Ma (Marzoli et al. 1999). Dickinson and Lawton (2001) reported that the Gondwanan Coahuila crustal block, which consists of the southern half of Texas and the northeastern corner of Mexico, was accreted onto Laurentia during the Permian along the Ouachita-Marathon suture. Farther south, and separated by the northwest-oriented Coahuila Transform fault, the Gondwanan Tampico, Del Sur, and Yucatan-Chiapas blocks form the eastern half of Mexico (shaded yellow, Fig. 1.4A). As Pangea began to break up,

the Mezcalera Plate was consumed by the advancing Farallon Plate west of the Gondwanan terranes and south of the Coahuila Transform (Dickinson and Lawton 2001).

Gulf of Mexico Rifting and Continental Extension From Oxfordian (early Late Jurassic) to Tithonian (Latest Jurassic), the Yucatan block appears to have rotated about 22° counterclockwise, while extensive salt was deposited on extended and attenuated continental crust, from the time of the second ridge jump in the central Atlantic to about 150 Ma (Fig. 1.4B). The block was rotated about a pole located presently at 24°N, 81.5°W (Hall and Najmuddin 1994). This rotation requires a north-south oriented transform fault offshore of eastern Mexico (Marton and Buffler 1994; Pindell 1994). We interpret the westward ridge jump in the central Atlantic at about 160 Ma to be linked to the clearing by the Florida shelf of the “Trinidad corner” on the north coast of South America. That change, which created space for the Gulf of Mexico to open, was coeval with the onset of Yucatan block rotation. Salt in the Gulf of Mexico generally can be divided into 2 large regions, the northern Gulf of Mexico salt basin and the Campeche salt basin (Fig. 1.1), which are interpreted to have formed contemporaneously (Winker and Buffler 1988; Salvador 1991; Angeles-Aquino et al. 1994; Marton and Buffler 1994; Pindell 1994). Using the distribution of Jurassic evaporite deposits as a geometrical constraint, White (1980) and White and Burke (1980) showed that the Yucatan block could be restored by clockwise rotation. They reasoned that the landward morphology of the southern Campeche salt margin, and the northern Gulf salt basin, represent rift valley walls that formed as the continental blocks separated. The original distribution of salt deposits in the Gulf of Mexico is probably closely related to the areal extent of attenuated continental crust. Prior to seafloor spreading between 160 Ma and 150 Ma, rotation of the Yucatan block and continental crustal extension allowed intermittent seawater influx that produced massive salt deposition. The lack of evidence for autochthonous salt (Peel et al. 1995; Hall 2001) beneath the Keathley Canyon anomaly probably means that the Keathley Canyon and Yucatan Parallel structures formed seaward boundaries for autochthonous Louann and Campeche salt as seafloor spreading continued until about 140 Ma. The Keathley Canyon structure is now hidden beneath a Plio-Pleistocene

Figure 1.4. Formation of Mexico, Gulf of Mexico, and the central Atlantic Ocean after Pangea breakup: (A) M40 (165.1 Ma), (B) M25 (154.1 Ma), and (C) M0 (124.6 Ma). Present western and northern coastlines of South America (west of TC) have been used for ease of geographic reference. Jurassic and Cretaceous coastlines in those regions, although poorly known, were certainly very different. North America (green) and South America (blue) are relative to Africa (black); South America-Africa closest-fit position for M40 and M25, and for M0 as South America drifted away from Africa, after Bird et al. (2005b); and present-day Yucatan and Chortis blocks relative to North America are light gray. One kilometer and 2 km isobaths, and estimated positions of abandoned central Atlantic seafloor-spreading centers (dotted lines), are plotted. Mexico crustal blocks (red), Ouachita-Marathon Suture (OM, magenta), and transform faults (heavy black) are modified after Dickinson and Lawton (2001). Bahama Islands (red) may overlie seamounts produced by the Central Atlantic Magmatic Province mantle plume. Yellow represents Gondwanan terranes. CP = Coahuila Platform, CT = Coahuila Transform, DS = Del Sur block, GS = Guerrero Superterrane, MC = Mesa Central Triassic subduction complex, MP = Mezcalara Plate, Tam = Tampico block, TC = “Trinidad corner,” TGLC = Tamaulipas-Golden Lane-Chiapas transform fault, and YB = Yucatan block. The heavy arrow, PPBC = the direction of Pre-Pangea Breakup Closure.

10 ~ Bird, Burke, Hall, and Casey

allochthonous salt nappe, but the Yucatan Parallel structure is clearly a boundary that separates the Campeche salt from the oceanic center of the basin.

Gulf of Mexico Seafloor Spreading By about 140 Ma, Tithonian (Latest Jurassic) to BerriasianValanginian (earliest Cretaceous), the Gulf of Mexico appears to have been completely formed after another 20° (42° total) of counterclockwise rotation by seafloor spreading (Fig. 1.4C). Crustal thicknesses from refraction data (Fig. 1.5) indicate typical passive margin conti-

nental thicknesses of over 20 km thinning to typical oceanic thicknesses of 4 to 8 km towards the center of the basin (Bird et al. 2005a). Crustal thicknesses under the Keathley Canyon and Yucatan Parallel gravity anomalies range from over 6.5 to 13 km and are similar to the thicknesses of crusts of seamounts produced by mantle plumes elsewhere in the world’s ocean basins (Bird et al. 2005a). Modeled cross sections (Fig. 1.5) constrained by seismic refraction and gravity data constructed for the Keathley Canyon and Yucatan Parallel structures indicate that the structures have similar dimensions to other hotspot structures (Bird et al. 2005a). The Keathley Canyon and Yucatan Parallel anomalies are similar in wavelength and

Figure 1.5. Seismic refraction control and modeled gravity cross-section locations in the western Gulf of Mexico. Bathymetry and topography contour interval=200 m, Keathley Canyon (KC), Yucatan Parallel (YP), and Tamaulipas-Golden Lane-Chiapas (TGLC) gravity anomaly outlines (dashed), 2.5-D model locations (A-A′, B-B′, C-C′, D-D′, and E-E′; Bird et al. 2005a), and seismic refraction information. Short solid-line segments coincide with seismic refraction profiles. Numbers expressed as fractions are generalized from literature sources and indicate depths in kilometers to the top and base of the crust; single numbers indicate depths to the top of crust only.

Tectonic Evolution ~ 11

amplitude to other anomalies produced by hotspot tracks such as the Galapagos Islands, New England Seamounts, Walvis Ridge, Rio Grande Rise, Ninetyeast Ridge, Hollister Ridge, Emperor Seamounts, and the Hawaiian Islands (Bird et al. 2005a). Thick and complex allochthonous salt over the Keathley Canyon structure masks its shape from seismic reflection data, but the existence of this large basement structure is clear in seismic refraction data over and near the structure (Ewing et al. 1960; Ibrahim et al. 1981; Ebeniro et al. 1988). Ewing et al. (1960, p. 4096) noted that a large ridge, composed of 5 km / s material, “separates the Sigsbee deep from the Gulf geosyncline.” Ebeniro et al. (1988) estimated the thickness of the Keathley Canyon structure to be 12 km and considered that the high-velocity layer associated with the top of the structure, beneath the MidCretaceous Unconformity, may be a basement structure. The narrow rectangular box in Figure 1.2 encloses trajectories for hotspot-referenced motion of North America for 140 Ma, 150 Ma, and 160 Ma (Morgan 1983). The trend of these trajectories and the overall trend of the Keathley Canyon anomaly are the same, indicating that if the Keathley Canyon structure is a hotspot track on the North American Plate, then it could have formed between 160 Ma and 140 Ma. Furthermore, the easternmost termination of the Yucatan Parallel structure also falls along the hotspot-referenced trajectories indicating no significant relative motion of the Yucatan with respect to North America after this time. We interpret the distinctive shapes of the Keathley Canyon and Yucatan Parallel anomalies to indicate that initially the velocities of the spreading center and hotspot track growth were similar, causing conjugate hotspot tracks to form on both the North American Plate and the Yucatan block (Figs. 1.6B, 1.6C). Later, the velocity of hotspot track growth increased relative to the velocity of the spreading center and the hotspot track continued to grow only on the Yucatan block. Therefore, although the Keathley Canyon track shows the relative motion between North America and the mantle plume, it records only part of the total opening history of the Gulf. The Yucatan Parallel track records the total rotation history during the seafloor-spreading phase of the evolution of the Gulf of Mexico (Figs. 1.6D, 1.6E). Reconstruction tracks from our opening scenario, with tracks calculated in 2.5° increments totaling 20° of seafloor spreading using an Euler pole from Hall and Nadjmuddin (1994) located about 100 km south of Key West at 24°N, 81.5°W (Fig. 1.6E), are superimposed on free air gravity anomalies in Figure 1.2.

As the Yucatan block rotated, a sheared margin was created along the east coast of central Mexico (Pindell 1985, 1994; Marton and Buffler 1994). Shear margins are ocean-continent transform or fracture zone boundaries and typically form after: (1) rupture of continental crust, rifting, and the formation of an intracontinental transform boundary, (2) the development of a seafloor-spreading center and a continent-oceanic transform boundary as the continental blocks slide past each other, and (3) thermal subsidence of the fracture-zone margin that separates oceanic from continental crust (Lorenzo 1997). Several examples of shear margins reveal that high-standing marginal ridges, rising 1 to 3 km over the abyssal seafloor and ranging from 50 to 100 km wide, form along the continental sides of these margins (Bird 2001). The formation of marginal ridges has been attributed to the absorption of heat from juxtaposed, thin (essentially zero at the spreading center), oceanic lithosphere as the ridge transform intersection moves past thick (over 100 km) continental lithosphere (Todd and Keen 1989; Lorenzo 1997). Marginal ridges can be topographic features such as the Côte d’Ivoire-Ghana marginal ridge, the Davie Ridge, and the Queen Charlotte Islands; or, depending on sedimentation rates, they can be completely buried by sediments such as in the southern Exmouth Plateau and the Agulhas-Diaz Ridges (Mascle et al. 1987; Mackie et al. 1989; Lorenzo et  al. 1991; Ben-Avraham et  al. 1997; Edwards et  al. 1997; Lorenzo and Wessel 1997). The Tamaulipas-Golden Lane-Chiapas anomaly in the Gulf of Mexico is not correlated with bathymetric relief and therefore must be attributed to a density contrast at depth. In both cases, marginal ridges produce prominent free air gravity anomaly highs that are similar in amplitude, wavelength, and orientation to the Tamaulipas-Golden Lane-Chiapas anomaly (free air gravity data derived from global satellite, Sandwell and Smith 1997). The anomalies are approximately 30 milligals (mGal) to 80 mGal in amplitude, 20 to 70 km in wavelength, and oriented parallel to bounding oceanic transforms or fracture zones. If the plume was active only during seafloor spreading, then the southern and eastern endpoints of the Yucatan Parallel structure, and the northwestern endpoint of the Keathley Canyon structure, are the southern, eastern, and northern limits of oceanic crust. The eastern flank of the Tamaulipas-Golden Lane-Chiapas structure (marginal ridge), along the east coast of central Mexico, defines the western limit of oceanic crust. The location of the oceanic-continental crustal boundary in the Gulf of

12 ~ Bird, Burke, Hall, and Casey

Figure 1.6. Hotspot referenced, seafloor-spreading phase of the opening of the Gulf of Mexico with a mantle plume. (A) Seafloor spreading is initiated over the mantle plume and the earliest formation of the hotspot tracks. (B) through (E) show the expected hotspot track geometry, and a schematic position of the spreading center (double lines), with four 5° steps. The seafloor spreading half-rate was roughly equal to the velocity of the North American plate over the mantle plume such that the plume remained beneath the spreading center for about 5 Myr (Figs. 1.6A–C) producing conjugate hotspot tracks (the Keathley Canyon [KC] and Yucatan Parallel [YP] tracks) on both the North American plate and the Yucatan block. Later (Figs. 1.6D, E), seafloor spreading slowed relative to hotspot growth and the mantle plume ended up beneath the Yucatan block (another 5 Myr). Rotations were calculated using an Euler pole (HN) described by Hall and Najmuddin (1994).

Mexico is interpreted along these areas as solid lines that are then connected by dashed lines in Figures 1.1 and 1.2.

Discussion The time required to span the distance from the northwesternmost end of the Keathley Canyon anomaly to the eastern end of the Yucatan Parallel anomaly, in the hotspot reference frame, is about 10 Myr (Morgan 1983), or nearly one-half the total time interval required for the Gulf of Mexico to open (Salvador 1987, 1991; Marton

and Buffler 1994). Since about 20° of clockwise rotation is needed to restore the western ends of the Keathley Canyon and Yucatan Parallel tracks, and this rotation must have occurred over the 10-Myr interval, then the rotation of about 20° should be roughly one-half the total rotation required to open the basin. These results, that the total time and rotation are approximately 20 Myr and 42° (Fig. 1.7), are consistent with evolutionary data presented by other workers. Exactly when this 20-Myr period occurred is difficult to determine, but stratigraphic relationships indicate that the basin must have been completely formed by ca. 140 Ma.

Tectonic Evolution ~ 13

Figure 1.7. Reconstruction of Gulf of Mexico, 20-Myr evolution of Yucatan motion, using rotation pole (HN) described by Hall and Najmuddin (1994). (A) Initial position about 160 Ma. Yucatan occupies what is the Gulf of Mexico basin now. Because the Yucatan was probably longer at that time, there was no gap between the peninsula and western Florida (Burke 1988). (B) 10 to 12 Myr: about 22° of rotation and continental crust extension. Seafloor spreading began at the end of this time when the plume became active. (C) After another 8 to 10 Myr, about 20° of rotation and seafloor spreading until the present geometry is achieved. Keathley Canyon (KC), Yucatan Parallel (YP), and Tamaulipas-Golden Lane-Chiapas (TGLC) gravity anomalies, Mid-Ocean Ridge (MOR), and ocean-continent boundary (OCB).

Our conclusion that seafloor spreading occurred between 150 Ma and 140 Ma implies that the Gulf of Mexico opened about 30 Myr after seafloor spreading began in the central Atlantic Ocean (Withjack et al. 1998). During that 20-Myr interval seafloor spreading between North and South America must also have been in progress. This allows us to distinguish several tectonic events in the evolution of North America and the Gulf of Mexico beginning with the breakup of Pangea (Table 1.1): onset of

rifting, salt deposition, onset of Yucatan rotation by continental extension, onset of seafloor spreading, and the end of seafloor spreading. As Pangea broke up, mantle plumes appear to have found older rifts and erupted before the plates drifted apart (Sleep 1997). The CAMP (200  Ma) and Karroo (183 Ma) plume eruptions preceded the breakup of North America, Australia-India-Antarctica, and Madagascar from Africa; the Bunbury member of the Kerguelen

14 ~ Bird, Burke, Hall, and Casey

Table 1.1. Chronology of tectonic events.

Conclusion

Time

Event

230 Ma

Pangea breakup began: collapse of the Appalachians and Ouachitas

230 to 164 Ma

Mesa Central Subduction complex began to form as the Mezcalera Plate is consumed by the Farallon Plate; Gondwanan crustal blocks south of the Coahuila Transform are displaced eastward; extension of the Coahuila block toward the southeast, and stretching of the Yucatan block

200 Ma

CAMP plume erupts

180 Ma

Seafloor spreading began in the Central Atlantic (Withjack et al. 1998)

170 Ma

Eastward ridge jump in the Central Atlantic (abandoning African lithosphere on the western flank)

160 Ma

Westward ridge jump in the Central Atlantic (abandoning North American lithosphere on the eastern flank)

~160 Ma

Yucatan block began to rotate away from North America, 24° counterclockwise continental extension

~150 Ma

Seafloor spreading in the Gulf of Mexico, 20° counterclockwise rotation of the Yucatan block

~140 Ma

Gulf of Mexico formation was complete

~126 Ma

South America began separating from Africa

120 Ma

Guerrero Superterrane was accreted onto western Mexico

A comparison of gravity anomalies over other hotspot tracks with the Keathley Canyon and Yucatan Parallel anomalies, and crustal structures of other hotspot tracks with 2-D modeling results, indicates that the Keathley Canyon and Yucatan Parallel anomalies are produced by deep-basement structures that are similar to the seamounts and seamount tracks created by mantle plumes. These structures are not continental fragments as indicated by their size, shape, and crustal structure. We suggest that these structures are hotspot tracks that were created by a single Late Jurassic mantle plume during the formation of the Gulf of Mexico basin (Bird et  al. 2001; Bird 2004). Another deep-basement structure (the Tamaulipas-Golden Lane-Chiapas marginal ridge) is consistent in size and shape with other marginal ridges around the world. The eastern flank of this ridge and the northern, eastern, and southern terminations of the hotspot tracks coincide with the oceanic-continental crustal boundary. Our proposed plate kinematic model and interpreted basement structures are consistent with established parameters including the pole of Yucatan block rotation, fracture zone, crustal types, the onset of rifting, early salt deposition, and deepwater marine sedimentation. Basin formation began with about 22° of counterclockwise rotation and continental extension (about 160 Ma to 150 Ma), which coincided with early salt deposition. Then another 20° of counterclockwise rotation and seafloor spreading coincided with the formation of hotspot tracks (about 150 Ma to 140 Ma).

(135 Ma) plume cluster preceded the breakup of India and Antarctica-Australia; and the Tristan (133 Ma) plume preceded the opening of the south Atlantic Ocean. Later the Marion, Deccan, and Iceland plumes (85 Ma, 65 Ma, and 60 Ma, respectively), preceded the breakup of the Seychelles from Madagascar, India from the Seychelles, and Greenland from the British Isles. After North America separated from the African-Arabian-South American continent (Residual Gondwana), Gondwanan terranes remained sutured to North America; that is, eastern Mexico, the Yucatan Peninsula, and the southern part of the Gulf States were contiguous from the Pacific to the Atlantic oceans. Only the Yucatan, which was surrounded on 3 sides by similar terranes, broke away to form the Gulf of Mexico. We consider the Gulf mantle plume to have played a similar role as other mantle plumes prior to continental breakup.

References Angeles-Aquino, F. J., J. Reyes-Nunez, J. M. Quezada-Muneton, and J. J. Meneses-Rocha. 1994. Tectonic evolution, structural styles, and oil habitat in Campeche Sound, Mexico. Gulf Coast Association of Geological Societies Transactions 44:53–62. Ben-Avraham, Z., C. J. H. Hartnady, and K. A. Kitchin. 1997. Structure and tectonics of the Agulhas–Falkland fracture zone. Tectonophysics 282:83–98. Bird, D. E. 2001. Shear margins: continent–ocean transform and fracture zone boundaries. The Leading Edge 20:150–59. ———. 2004. Jurassic tectonics of the Gulf of Mexico and central Atlantic Ocean. Ph.D. diss., University of Houston, Houston, Tex. 173 pp. Bird, D. E., K. Burke, S. A. Hall, and J. F. Casey. 2005a. Gulf of

Tectonic Evolution ~ 15

Mexico tectonic history: hotspot tracks, crustal boundaries, and early salt distribution. American Association of Petroleum Geologists Bulletin 89:311–28. Bird, D. E., S. A., Hall, K. Burke, and J. F. Casey. 2005b. Late Jurassic–Early Cretaceous tectonic reconstructions of the Central and South Atlantic Oceans. Eos, Transactions, American Geophysical Union, Joint Assembly Supplement 86:JA508-JA509. Bird, D. E., S. A., Hall, J. F. Casey, and K. Burke. 2001. Geophysical evidence for a possible late Jurassic mantle plume in the Gulf of Mexico. Eos, Transactions, American Geophysical Union, Fall Meeting Supplement 82:F1185. Brozena, J. M. 1986. Temporal and spatial variability of sea-floor spreading processes in the northern South Atlantic. Journal of Geophysical Research. 91:497–510. Buffler, R. T., and W. A. Thomas. 1994. Crustal structure and evolution of the southeastern margin of North America and the Gulf of Mexico. Pp. 219–64 in R. C. Speed (ed.), Phanerozoic Evolution of North American Continent–Ocean Transitions. DNAG Continent-Ocean Transect Volume. Boulder, Colo.: Geological Society of America. Burke, K. 1988. Tectonic evolution of the Caribbean. Annual Review of Earth and Planetary Sciences 16:201–30. Burke, K., D. S. Macgregor, and N. R. Cameron. 2003. Africa’s petroleum systems: four tectonic “Aces” in the past 600 million years. Pp. 21–60 in T. J. Arthur, D. S. MacGregor, and N. R. Cameron (eds.), Petroleum Geology of Africa: New Themes and Developing Technologies. Special Publication 207. London: Geological Society. Christenson, G. 1990. The Florida lineament. Transactions, Gulf Coast Association of Geological Societies 40:99–115. Dewey, J. F. 1988. Extensional collapse of orogens. Tectonics 7:1123–139. Dickinson, W. R., and T. F. Lawton. 2001. Carboniferous to Cretaceous assembly and fragmentation of Mexico. Geological Society of America Bulletin 113:1142–160. Dietz, R. S. 1973. Morphologic fits of North America / Africa and Gondwana: a review. Pp. 865–72 in D. H. Tarling and S. K. Runcorn (eds.), Implications of Continental Drift to the Earth Sciences, vol. 2. New York:Academic Press. Dunbar, J. A., and D. S. Sawyer. 1987. Implications of continental crust extension for plate reconstruction: an example from the Gulf of Mexico. Tectonics 6:739–55. Ebeniro, J. O., Y. Nakamura, D. S. Sawyer, and W. P. O’Brien Jr. 1988. Sedimentary and crustal structure of the northwestern Gulf of Mexico. Journal of Geophysical Research 93:9075–92. Edwards, R. A., R. B. Whitmarsh, and R. A. Scrutton. 1997. Synthesis of the crustal structure of the transform continen-

tal margin off Ghana, northern Gulf of Guinea. Geo-Marine Letters 17:12–20. Ewing, J., J. Antoine, and M. Ewing. 1960. Geophysical measurements in the western Caribbean Sea and in the Gulf of Mexico. Journal of Geophysical Research 65:4087–126. Gose, W. A., R. C. Belcher, and G. R. Scott. 1982. Paleomagnetic results from northeastern Mexico: evidence for large Mesozoic rotations. Geology 10:50–54. Gradstein, F., J. Ogg, and A. Smith. 2004. A Geologic Time Scale 2004. Cambridge, U.K.: Cambridge University Press. 589 pp. Hall, S. 2001. The development of large structures in the deepwater northern Gulf of Mexico. Houston Geological Society Bulletin 43:8:20–23. Hall, S. A., and I. J. Najmuddin. 1994. Constraints on the tectonic development of the eastern Gulf of Mexico provided by magnetic anomaly data. Journal of Geophysical Research 99:7161–175. Humphris, C. C. Jr. 1979. Salt movement on continental slope, northern Gulf of Mexico. American Association of Petroleum Geologists Bulletin 63:782–98. Ibrahim, A. K., J. Carye, G. Latham, and R. T. Buffler. 1981. Crustal structure in Gulf of Mexico from OBS refraction and multichannel reflection data. American Association of Petroleum Geologists Bulletin 65:1207–229. Lorenzo, J. M. 1997. Sheared continent–ocean margins: an overview. Geo-Marine Letters 17:1–3. Lorenzo, J. M., J. C. Mutter, R. L. Larson, and Northwest Australia Study Group. 1991. Development of the continent—ocean transform boundary of the southern Exmouth Plateau. Geology 19:843–46. Lorenzo, J. M., and P. Wessel. 1997. Flexure across a continent– ocean fracture zone: the northern Falkland /Malvinas Plateau, South Atlantic. Geo-Marine Letters 17:110–18. Luhr, J. F., S. A. Nelson, J. F. Allan, and S. E. Carmichael. 1986. Active rifting in southwestern Mexico: manifestations of an incipient eastward spreading-ridge jump. Geology 13:54–57. Mackie, D. J., R. M. Clowes, S. A. Dehler, R. M. Ellis, and P. Morel-À-L’Huissier. 1989. The Queen Charlotte Islands refraction project. Part II. Structural model for transition from Pacific plate to North American plate. Canadian Journal of Earth Sciences 26:1713–725. Mammerickx, J., D. F. Naar, and R. L. Tyce. 1988. The Mathematician paleoplate. Journal of Geophysical Research 93:3025–40. Mammerickx, J., and D. Sandwell. 1986. Rifting of old oceanic lithosphere. Journal of Geophysical Research 91:1975–988. Martin, R. G. 1980. Distribution of Salt Structures in the Gulf of Mexico. USGS Miscellaneous Field Studies Map MF-1213. Boulder, Colo.: U.S. Geological Survey.

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Marton, G., and R. T. Buffler. 1994. Jurassic reconstruction of the Gulf of Mexico Basin. International Geology Review 36:545–86. Marzoli, A., P. R. Renne, E. M. Piccirillo, M. Ernesto, G. Bellieni, and A. De Min. 1999. Extensive 200-million-year-old continental flood basalts of the Central Atlantic Magmatic Province. Science 284:616–18. Mascle, J., D. Mougenot, E. Blarez, M. Marinho, and P. Virlogeux. 1987. African transform continental margins: examples from Guinea, the Ivory Coast and Mozambique. Geological Journal 22:537–61. Molina-Garza, R. S., R. Van der Voo, and J. Urrutia-Fucugauchi. 1992. Paleomagnetism of the Chiapas Massif, southern Mexico: evidence for rotation of the Maya Block and implications for the opening of the Gulf of Mexico. Geological Society of America Bulletin 104:1156–168. Morgan, W. J. 1983. Hotspot tracks and the early rifting of the Atlantic. Tectonophysics 94:123–39. Morton, J. L., and R. D. Ballard. 1986. East Pacific Rise at lat 19°S: evidence for a recent ridge jump. Geology 14:111–14. Muller, R. D., W. R. Roest, J.-Y. Royer, L. M. Gahagan, and J. G. Sclater. 1997. Digital isochrons of the world’s ocean floor. Journal of Geophysical Research 102:3211–214. Peel, E. J., J. R. Hossack, and C. J. Travis. 1995. Genetic structural provinces and salt tectonics of the Cenozoic offshore U.S. Gulf of Mexico: a preliminary analysis. Pp. 153–75 in M. P. A. Jackson, D. G. Roberts, and S. Snelson (eds.), Salt Tectonics: A Global Perspective. Memoir 65. Boulder, Colo.: American Association of Petroleum Geologists. Pindell, J. L. 1985. Alleghenian reconstruction and subsequent evolution of the Gulf of Mexico, Bahamas and proto-Caribbean. Tectonics 4:1–39. ———. 1994. Evolution of the Gulf of Mexico and the Caribbean. Pp. 13–39 in S. K. Donovan, and T. A. Jackson (eds.), Caribbean Geology: An Introduction. Kingston, Jamaica: The University of the West Indies Publishers Association. Pindell, J., and J. F. Dewey. 1982. Permo-Triassic reconstruction of western Pangea and the evolution of the Gulf of Mexico / Caribbean region. Tectonics 1:179–211. Ross, M. I., and C. R. Scotese. 1988. A hierarchical tectonic

model of the Gulf of Mexico and Caribbean region. Tectonophysics 155:139–68. Salvador, A. 1987. Late Triassic–Jurassic paleogeography and origin of Gulf of Mexico Basin. American Association of Petroleum Geologists Bulletin 71:419–51. ———. 1991. Origin and development of the Gulf of Mexico basin. Pp. 389–444 in A. Salvador (ed.), The Geology of North America, Volume J, The Gulf of Mexico Basin, Boulder, Colo.: Geological Society of America. Sandwell, D. T., and W. H. F. Smith. 1997. Marine gravity anomaly from Geosat and ERS 1 satellite altimetry. Journal of Geophysical Research 102:10039–54. Schouten, H., and K. D. Klitgord. 1994. Mechanistic solutions to the opening of the Gulf of Mexico. Geology 22:507–10. Shepherd, A. V. 1983. A study of the magnetic anomalies in the eastern Gulf of Mexico. M.S. thesis, University of Houston, Houston, Tex. 197 pp. Sleep, N. H. 1997. Lateral flow and ponding of starting plume material. Journal of Geophysical Research 102:10001–12. Todd, B. J., and C. E. Keen. 1989. Temperature effects and their geological consequences at transform margins. Canadian Journal of Earth Science 26:2591–603. Vogt, P. R., C. N. Anderson, and D. R. Bracey. 1971. Mesozoic magnetic anomalies, sea-floor spreading, and geomagnetic reversals in the southwestern North Atlantic. Journal of Geophysical Research 76:4796–823. White, G. W. 1980. Permian–Triassic continental reconstruction of the Gulf of Mexico–Caribbean area. Nature 283:823–26. White, G. W., and K. C. Burke. 1980. Outline of the tectonic evolution of the Gulf of Mexico and Caribbean region. Houston Geological Society Bulletin 22:10:8–13. Winker, C. D., and R. T. Buffler. 1988. Paleogeographic evolution of early deep-water Gulf of Mexico and margins, Jurassic to Middle Cretaceous (Comanchean). American Association of Petroleum Geologists Bulletin 72:318–46. Withjack, M. O., R. W. Schlische, and P. E. Olsen. 1998. Diachronous rifting, drifting, and inversion on the passive margin of central eastern North America: an analog for other passive margins. American Association of Petroleum Geologists Bulletin 82:817–35.

␥2

Geology of the Florida Platform Pre-Mesozoic to Recent Thomas M. Scott

The Florida Platform is delimited by the 200 m (600 ft) isobath at the shelf break to the approximate location of the Paleozoic suture beneath southern Georgia and Alabama (Fig. 2.1). The Suwannee–Wiggins Suture (Thomas et al. 1989) is the proposed location where terranes with African affinities are welded to the North American Plate (Chowns and Williams 1983; McBride and Nelson 1988; Woods et  al. 1991). The basement rocks of the Florida Platform are a fragment of the African Plate that remained attached to the North American Plate when rifting occurred in the Jurassic and range in age from late Precambrian-early Cambrian to mid-Jurassic (Barnett 1975). Excellent reviews of the geology of the basement are provided by Smith (1982), Arthur (1988), Smith and Lord (1997), and Heatherington and Mueller (1997). Barnett (1975) provided a structure contour map of the sub-Zuni surface. This surface equates to what is now recognized as pre-Middle Jurassic. Barnett’s interpretation of the basement surface has it occurring as shallow as approximately 915 m (3000 ft) below mean sea level (msl) in central-northern peninsular Florida. The basement surface dips west and southwest toward the Gulf of Mexico basin, to the south into the South Florida basin, and to the east into the Atlantic basin. The basement surface reaches depths of more than 5180 m (17,000 ft) below msl in southern Florida (Barnett 1975). The platform, deposited unconformably on top of the basement, is constructed of Middle Jurassic to Holocene

evaporite, carbonate, and siliciclastic sediments deposited on a relatively stable, passive margin of the North American Plate. The age assignments for the Middle Jurassic to Holocene formations are, at times, tentative propositions due to limited, or lack of, paleontological evidence in some formations. The age determinations for some of the younger units, for example the Pliocene Tamiami Formation, are based on a vast amount of paleontological evidence. This, in part, is responsible for differing interpretations of when, where, and how much sediment was deposited across the platform (see and compare Salvador [1991b] and Randazzo [1997]).

Structure The Florida Platform has been a relatively stable portion of the trailing edge of the North American Plate since the mid-Jurassic. Winston (1991) stated that the Mesozoic and Cenozoic structural movement on the Florida– Bahama Platform was entirely negative. Florida’s arches, or structural highs, were not formed by uplift but as the result of subsiding more slowly than the flanking basins. However, faulting of the basement rocks created many of the structural features recognized on the pre–midJurassic surface (Barnett 1975; Smith and Lord 1997). Faults disrupting the Upper Jurassic sediments have been identified in northwestern Florida; some displacements exceed 305 m (1000 ft) (Lloyd 1989). Miller (1986) rec-

17

18 ~ Scott

80°0′0″W

85°0′0″W

80°0′0″W

85°0′0″W

80°0′0″W

30°0′0″N

30°0′0″N

Georgia Channel System boundary

85°0′0″W

25°0′0″N

25°0′0″N

25°0′0″N

25°0′0″N

30°0′0″N

30°0′0″N

Ap

ala

chic ola

85°0′0″W

80°0′0″W

Figure 2.1. Limits of the Florida Platform.

Figure 2.2. Structures affecting the Mesozoic and early Cenozoic deposits (modified after Lloyd 1997).

ognized a number of known or suspected Cenozoic faults that affect the Floridan aquifer system. Duncan et  al. (1994) identified faulting in the Lower to Middle Eocene Oldsmar Formation. A number of hydrogeologic and geomorphic investigations have proposed the existence of faults (Wyrick 1960; Leve 1966; Lichtler et al. 1968; Pirkle 1970; White 1970). The faults in the Cenozoic section have very limited displacement, generally less than 30.5 m (100 ft) and are difficult to identify due to limited displacement, well control, few “marker” beds, erosional disconformities, and karstification. Little has been said concerning folding of post–midJurassic sediments on the Florida Platform. Missimer and Maliva (2004) believe that folding is more widespread on the Florida Platform than is presently recognized due to the limited amount of detailed subsurface data. They recognized folding with associated fracturing and faulting in the sediments of the intermediate (Miocene–Pliocene sediments) and Floridan aquifer systems (Eocene–Oligocene sediments) on the southern portion of the platform. They postulated that the interaction of the Caribbean and

North American Plates in the Late Miocene to Pliocene produced the folds, fractures, and faults. The oldest features recognized as affecting deposition of post–mid-Jurassic sediments on the platform are expressed on the pre–mid-Jurassic surface (Arthur 1988). The Mesozoic structural features affecting deposition of sediments include a series of basins or embayments and arches (Fig. 2.2). Some of these features affected deposition into the mid-Cenozoic (for example, the South Florida basin; Scott 1988). Other features affected the deposition into the late Cenozoic (for example, the Apalachicola Embayment; Schmidt 1984). The Peninsular Arch affected deposition from the Jurassic through the Cretaceous and was intermittently positive during the Cenozoic (Miller 1986). The Cenozoic structural features affecting deposition are shown in Figure 2.3. One of the more interesting structural features of the Florida Platform is a southwest-to-northeast trending low that has affected deposition from the mid-Jurassic until at least the Middle Miocene. Some portions of the feature continued to affect deposition through the Pleisto-

Geology of the Florida Platform ~ 19

Depositional Environments

80°0′0″W

30°0′0″N

30°0′0″N

85°0′0″W

Georgia Channel System boundary Axis and dip direction along geologic feature topographic high or positive feature

25°0′0″N

25°0′0″N

topographic low or basin

85°0′0″W

80°0′0″W

Figure 2.3. Structures affecting the post- early Cenozoic deposits (modified after Scott 1988).

cene. This feature has an extended list of names that have been applied to all or parts of it. An excellent review of the names applied to the feature was presented by Schmidt (1984) and Huddlestun (1993). However, Georgia Channel System is the name that has been applied to the entire sequence (Huddlestun 1993) (Figs. 2.2, 2.3). The Georgia Channel System had its origin in the formation of the South Georgia Rift in the Triassic– Jurassic(?) (Huddlestun 1993). From the Late Cretaceous through the Paleocene, this area was the boundary between carbonate deposition to the south and siliciclastic deposition to the north. By the Eocene, the Appalachian Mountains had been highly eroded leaving relatively low hills and significantly reduced siliciclastic sediment transport via streams and rivers. In the Eocene and Oligocene, as the result of a greatly reduced siliciclastic supply, carbonate deposition extended across the Georgia Channel System. The channel system was then infilled by predominantly siliciclastic sediments in the Late Oligocene to the Early Miocene in response to uplift in the Appalachians (Scott 1988).

The initial depositional environments affecting the Florida Platform were restricted environments allowing for intense evaporation and the development of evaporites in limited areas. As the Gulf continued to expand and sea levels rose, siliciclastic and carbonate depositional environments began to cover more of the platform. Continued sea-level rise through the Cretaceous eventually covered the exposed land area in northern Florida. The Florida Platform sediments were deposited in a complex interplay of siliciclastic, carbonate, and evaporite facies as a result of sea-level fluctuations (Randazzo 1997). Siliciclastic deposition predominated on the northern part of the platform while carbonate and evaporite sediments formed to the south (Randazzo 1997). In the early Cenozoic (Paleogene), the siliciclastic sediment supply was limited due to the highlands of the Appalachian trend having been reduced by erosion, and carbonate deposition expanded to cover the entire Florida Platform and beyond by the Oligocene. The carbonate platform, which began as a rimmed shelf in the Jurassic, evolved to a carbonate ramp sequence by the early Cenozoic (Randazzo 1997; Winston 1991). Subsequent to the maximum development of the carbonate platform, uplift occurred in the Appalachians providing a renewed supply of siliciclastic sediments (Scott 1988; Brewster-Wingard et al. 1997). This influx of siliciclastic sediments in the Neogene replaced most carbonate deposition on the Florida Platform by the mid-Pliocene. As sea level rose in the late Pleistocene, there was a decrease in siliciclastic sedimentation and carbonate deposition increased on the southern Florida Platform. The interplay of the carbonate and siliciclastic sediments with fluctuating sea level and changing climate created complex depositional environments (Scott 1988; Missimer 2002). The interaction of the carbonates and siliciclastics on the Florida Platform has been investigated and discussed by a number of authors (Warzeski et al. 1996; Cunningham et al. 1998; Guertin 1998; Guertin et al. 2000; Missimer et al. 2000; Missimer 2001, 2002; Cunningham et al. 2003).

Stratigraphy Stratigraphically, Florida is composed of pre-Mesozoic sedimentary, igneous, and metamorphic rocks overlain by Mesozoic and Cenozoic sedimentary rocks. The

20 ~ Scott

Mesozoic sediments consist predominantly of siliciclastics except in central and southern Florida where carbonates predominate. In the Cenozoic, the Paleogene sediments are predominantly carbonates with some mixed carbonate-siliciclastic sediments. The Neogene and Quaternary sediments are predominantly siliciclastics (Braunstien et al. 1988). The pre-Mesozoic rocks occur nearest to the land surface in northern peninsular Florida. These rocks dip deeper in the subsurface to the south under the exposed portion of the Florida Platform, to the east under the Atlantic Ocean and west into the Gulf of Mexico (Puri and Vernon 1964). Consequently, the Mesozoic and Cenozoic sediments thicken in these areas exceeding 13,000 feet thick in southern Florida.

(early Paleozoic) (Dallmeyer et al. 1987; Smith and Lord 1997). Dallmeyer et al. (1987) recognized that these rocks, the Osceola Granite, were part of a complex that is also found in northwestern Africa. A metamorphic sequence, located on the southern flank of the Osceola Granite, indicates metamorphism was associated with the emplacement of the granite pluton. Sedimentary rocks are found in 2 areas of the basement, a small area in the panhandle near the junction of Alabama, Florida, and Georgia, and in the northern peninsula north of a line connecting Tampa Bay in the southwest and a point between St. Augustine and Jacksonville in the northeast (Jones 1997; Smith and Lord 1997). With the exception of sediments encountered in a few wells, the sandstone, siltstone, and shale are usually sparsely fossiliferous. Ages derived from the fossil assemblages range from Early Ordovician to Middle Devonian (Jones 1997). Opdyke et  al. (1987) recognized a pre-Mesozoic shale in northern Florida that exhibited low-grade metamorphism.

Pre-Mesozoic The pre-Mesozoic (Proterozoic and Paleozoic) framework of the Florida basement is composed of igneous, sedimentary, and metamorphic rocks (Fig. 2.4). These rocks have been penetrated by oil exploration boreholes. A number of researchers have investigated the preMesozoic rocks including Smith (1982), Chowns and Williams (1983), Dallmeyer et al. (1987), Arthur (1988), and others. Refer to Smith and Lord (1997) for a summary of the research on the basement rocks. The granitic igneous rocks, which occur in east-central Florida, have been dated at approximately 530 million years old

290

Unknown

Unknown

354

Unknown

Unnamed Sediments

417

Unknown

Unnamed Sediments 443

Unnamed Sediments 490

Unnamed Volcanic and Plutonic Complex

Unnamed Volcanic and Plutonic Complex

Unnamed Metamorphic Complex

Unnamed Granitic Complex 543

ERA T H EM Pal eozoi c

248

Pr ot er ozoi c

AGE (Ma)

S. FLORIDA

SY ST EM

N. FLORIDA

Mesozoic sediments on the Florida Platform were deposited in response to the separation of plates beginning in the Triassic. Subsequent to the breakup of the plates, marine sedimentation began and remained the dominant depositional type for much of the geologic history of the Platform.

Pr ecambr i am Cambr i an Or dovi ci an Si l ur i an Devoni an Car boni f er ous Per mi an

PANHANDLE

Mesozoic

Figure 2.4. Paleozoic stratigraphic columns (modified after Braunstein et al. 1988).

Geology of the Florida Platform ~ 21

The Gulf of Mexico basin began to form in the Late Triassic as rifting began to separate the lithospheric plates (Salvador 1991a). The first post-rifting sediments deposited on the Florida Platform were upper Middle Jurassic evaporites in the Apalachicola Embayment and the Conecuh Embayment (Fig. 2.2). These were deposited in very limited portions of the northwestern Florida Platform (Salvador 1991b; Randazzo 1997). Deltaic to shallow-marine siliciclastics, carbonates, and evaporites were depos-

PANHANDLE

N. FLORIDA

Cretaceous By the beginning of the Cretaceous, a limited portion of the northern Florida Peninsula remained above sea level

S. FLORIDA

SY ST EM

Jurassic

65

Lawson Formation

Selma Group

Rebecca Shoals Dolomite

Lawson Formation

Pine Key Formation

Pine Key Formation Orlando Sound Dolomite

Eutaw Formation

Atkinson Formation

Atkinson Formation Cr et aceous

Tuscaloosa Group

Naples Bay Group Unnamed Sediments

Ocean Reef Group

Sunniland Fm

Glades Group

M esozoi c

Sligo-Hosston Formation

Marquesas Supergroup

Big Cypress Group

Unnamed Sediments Unnamed Sediments

ERA T H EM

Triassic rifting associated with the breakup of Pangea and the formation of the Atlantic Ocean created the South Georgia basin (Rift) (Fig. 2.2). Triassic red beds, the Newark Group, and Eagle Mills Formation (Braunstein et al. 1988) (Fig. 2.5), filled the rift system. Basalt and diabase (tholeiites), with an average age of 192 million years (Arthur 1988), have been encountered in a number of boreholes. These rocks were emplaced or occurred as flows in response to the continued separation of the plates (Arthur 1988).

ited on the northwestern Florida Platform during the Late Jurassic (Salvador 1991b). These sediments contain important petroleum-producing horizons, including the Norphlet Sandstone and the Smackover Formation (carbonates) (Braunstein et al. 1988) (Fig. 2.5) that were discovered between 1970 and 1986 (Lloyd 1997). In southern Florida, Upper Jurassic siliciclastics were followed by carbonates and evaporites deposited on an unnamed Upper Triassic to Upper Jurassic volcanic complex (Braunstein et  al. 1988). These sediments occur below important petroleum producing horizons in the South Florida basin (Applegate et al. 1981). Throughout the mid-Jurassic to the beginning of the Cretaceous, sea levels rose, progressively reducing the exposed portion of the Florida Platform (Salvador 1991b; Randazzo 1997). The thickness of post–mid-Jurassic to Cretaceous sediments in the northwestern Florida Platform exceeds 1000 m (3500 ft) (Randazzo 1997). In the southern part of the platform, the thickness may exceed 915 m (3000 ft) (Winston 1987).

AGE (Ma)

Triassic

Pumpkin Bay Formation Bone Island Formation 144

Ft. Pierce Formation

Cotton Valley Formation

Wood River Formation

Haynesville Formation Jur assi c

Cotton Valley Formation Smackover Formation Norphlet SS

Unnamed Volcanic Complex

Louann Salt / Werner Anhydrite Diabase

206

Eagle Mills Formation

Newark Group

T r i assi c

Diabase

248

Figure 2.5. Mesozoic stratigraphic columns (modified after Braunstein et al. 1988).

22 ~ Scott

(McFarlan and Menes 1991). As sea level rose through the Early Cretaceous, more of the platform was submerged (McFarlan and Menes 1991; Randazzo 1997). Deposition in the northwestern Florida Platform was dominated by marine and non-marine siliciclastics. Carbonates and evaporites covered the southern portion of the platform (McFarlan and Menes 1991; Winston 1987, 1991). During the Lower Cretaceous, carbonates and evaporites of the Ocean Reef Group, Sunniland Formation (Fig. 2.5), and associated units were deposited. The Sunniland sediments became the reservoir rocks for Florida’s first oil discovery (1943) (Lloyd 1997). The thickness of the Lower Cretaceous sediments reaches more than 1830 m (6000 ft) on the northwestern and 2740 m (9000 ft) on the southern portions of the platform (Randazzo 1997). In the early portion of the Late Cretaceous, sediments in the northern portion of the Florida Platform continued to be dominated by siliciclastics, while carbonates were being deposited in southern Florida (Sohl et  al. 1991; Winston 1991). By the mid-Late Cretaceous, carbonates, including chalk, with limited siliciclastics were deposited over the entire Florida Platform (Sohl et al. 1991). The Upper Cretaceous sediments are more than 915 m (3000 ft) thick in northwestern and southern Florida (Randazzo 1997). At the end of the Cretaceous, a large bolide (meteorite, asteroid, or comet) collided with Earth in the Gulf of Mexico–Caribbean region (Hildebrand et al. 1991). The bolide impacted at an oblique angle, spreading ejecta to the north and west (Schultz 1996). It is thought that 100 to 300-m (330 to nearly 1000 ft) high tsunamis (Bourgeois et  al. 1988; Matsui et  al. 1999) spread across the Gulf of Mexico (Kring 2000). Discussions with a number of geologists investigating the Chicxulub impact suggest that the Florida Platform should have been influenced by the event (Chicxulub planning meeting–Group on Mesozoic–Cenozoic stratigraphy and the Cretaceous– Tertiary (KT) boundary, Puerto Vallarta, Mexico, 1993). However, no evidence of the impact or tsunamis has been discovered on the Florida Platform to date. The lack of cores across the KT boundary, the limited number and wide distribution of wells penetrating the KT boundary, and the general poor quality of the cuttings from the wells hinder the search for evidence.

Cenozoic Carbonate sedimentation dominated during the Paleogene and into the earliest Neogene on much of the Florida

Platform. A significant change in sedimentation occurred in the early Neogene. Siliciclastic sediments began to replace carbonates as the dominant sediment. Paleogene Carbonate–evaporite deposition dominated much of the Florida Platform during the Paleocene (Miller 1986). The carbonate–evaporite sediments graded to the northwest into shallow marine fine-grained siliciclastic sediments across the Georgia Channel System. The main carbonate-producing area was interpreted to be rimmed by a reef system creating the restricted environment necessary for evaporite deposition (Winston 1991). The Paleocene sediments cover the entire Florida Platform and have a maximum thickness of more than 670  m (2200 ft). The thick anhydrite beds in the Cedar Keys Formation (Fig. 2.6) form the regionally extensive lower confining bed of the Floridan aquifer system (Miller 1986, 1997). The evaporite content of the Lower to Middle Eocene sediments declined in response to sea-level rise and resulted in the development of a more open, carbonateramp depositional system on the platform. Evaporites occur primarily as pore fill (Miller 1986). The carbonate sediments grade into siliciclastic sediments in the Georgia Channel System (Miller 1986). The Lower to Middle Eocene sediments cover the entire platform, ranging to maximum thickness of more than 945 m (3100 ft). Middle Eocene carbonates (Avon Park Formation) are the oldest sediments exposed on the platform (Scott et al. 2001). These sediments crop out on the crest of the Ocala Platform (Fig. 2.3). The Lower to Middle Eocene limestone and dolostone, in part, form the lower portion of the Floridan aquifer system while, in some areas, these sediments are part of the lower confining bed of the aquifer system (Miller 1986, 1997). Carbonate deposition covered virtually the entire Florida Platform in the Late Eocene. Carbonates were deposited to the north of the Georgia Channel System nearly to the Fall Line (limit of Cretaceous overlap), beyond the limits of the Florida Platform (Fig. 2.1). The carbonate ramp was well developed and evaporites have not been found in the limestone or dolostone. The carbonates grade into siliciclastics on the northwesternmost part of the platform. Upper Eocene carbonates range in thickness to more than 213 m (700 ft) but, due to erosion, are absent in several areas of the platform (Miller 1986; Scott 1992, 2001). In a large area on the southern part of the platform,

Undifferentiated Sediments Undifferentiated Holocene - Pleistocene Sediments

Anastasia Miami Key Largo Formation Limestone Limestone Fort Thompson Formation Bermont Beds

Intracoastal Formation

Miccosukee Formation

Jackson Bluff Formation

Nashua Formation

1.8

Caloosahatchee Formation Tamiami Formation

Choctawhatchee Formati on

ER A T H E M

SERI ES

5.3

Marks Head Formation

Penney Farms Formation

Wilcox Group

Suwannee Limestone Ocala Limestone

33.7

Ocala Limestone

Avon Park Formation

Avon Park Formation

Oldsmar Formation

Oldsmar Formation

Cedar Keys Formation

Cedar Keys Formation

54.8

Rebecca Shoals Dolomite 65

the Upper Eocene sediments are absent, probably due to erosion by currents similar to episodes identified in the Oligocene to Pliocene in this region (Guertin et al. 2000). On the areas of the platform where the Oligocene carbonates are absent, the Upper Eocene carbonates form the upper Floridan aquifer system (Miller 1986, 1997). Lower Oligocene carbonate deposition occurred as far updip as did the Upper Eocene deposition. The carbonates grade into siliciclastics on the northwesternmost part of the platform. Very minor amounts of siliciclastics are incorporated in these carbonates. However, beds of fine quartz sand occur in the Lower Oligocene of southern Florida (Missimer 2002). Whether or not the carbonate deposition covered the platform is open to conjecture. The Lower Oligocene sediments range in thickness to more than 213 m (700 ft) but are absent over large portions of the platform (Miller 1986; Scott 1992, 2001). These sediments are missing due to nondeposition or erosion, or both, in a large area on the eastern flank of the Ocala Platform in an area referred to as the paleo-Orange Island

Cenozoi c Pal eogene

Ol i gocene

Nocatee Member

Eocene

Suwannee Limestone Vicksburg Group

Claiborne Group

M i ocene 23.8

Chickasawhay Formation

Ocala Limestone

Arcadia Formation Tampa Member

Pal eocene

Chipola Formation Chattahoochee Formation

Peace River Formation

Hawthorn Group

A l um Bl uf f Formati on

Shoal Ri ver Formati on

Coosawhatchie & Statesville Formations

Neogene

Bone Valley Member

Oak Grove Sands

I ntracoastal Formati on Bruce Creek L i mestone

Pensacol a Cl ay

Cypresshead Formation

Hawthorn Group

Citronelle Formation

10,000 yrs

Pl i ocene

Undifferentiated Holocene - Pleistocene Sediments

SY ST EM

S. FLORIDA

Quar t er nar y

N. FLORIDA

Pl ei st ocene H ol ocene

PANHANDLE

AGE (Ma)

Geology of the Florida Platform ~ 23

Figure 2.6. Cenozoic stratigraphic columns (modified after Braunstein et al. 1988).

(Bryan 1991). Where the Lower Oligocene sediments are present, they form the upper portion of the Floridan aquifer system (Miller 1986, 1997). Chert (silicified limestone) occurs from the upper portion of the Middle Eocene carbonates through the Lower Oligocene carbonates. The chert formed as the result of the weathering of the overlying clay-rich Miocene sediments that covered the platform (Scott 1988). Weathering of the clays releases large amounts of silica into the groundwater and, in the appropriate geochemical environment, replaces limestone. Groundwater beneath the present-day erosional scarp near Lake City in northern Florida is supersaturated with respect to Opal-CT and slightly saturated with respect to quartz due to weathering of the clays (S. B. Upchurch, personal communication 2005). Fossils including foraminifera and corals are often preserved in the chert. Sea-level lowering in the Late Oligocene restricted deposition to portions of southern and northwestern Florida (Missimer 2002). Although absent over much of

24 ~ Scott

A

Figure 2.7. (A) Cross sections showing the shallow subsurface and surface distribution of Paleogene, Neogene, and Quaternary lithostratigraphic units (Scott et al. 2001), and (B) locations of cross sections (next page).

the platform, these sediments may exceed 135 m (440 ft) in thickness (Braunstein et al. 1988). The stratigraphic section in southern Florida may represent the most complete Upper Oligocene section in the southeastern United States (Brewster-Wingard et  al. 1997). In very limited

areas, the Upper Oligocene carbonates may form the top of the Floridan aquifer system (Miller 1986, 1997). Cross sections showing the distribution of the Paleogene sediments are shown in Figure 2.7. A generalized geologic map of Florida is shown in Figure 2.8. The Paleo-

Geology of the Florida Platform ~ 25

B

Figure 2.7. (continued)

gene lithostratigraphic units occurring in the surface and shallow subsurface of the panhandle, northern, and southern portions of Florida are shown in Figure 2.9. Neogene–Quaternary Significant depositional changes occurred in the latest Paleogene–earliest Neogene. Several factors were responsible for the changes including epeirogeny in the Appalachians that took a highly eroded and reduced mountain range and uplifted it (Stuckey 1965; Schlee et al. 1988). The rejuvenated mountain range again became a source of sediment due to increased erosion, and the siliciclastic sediments were transported by streams and rivers; marine currents transported the sediment southward onto the Florida Platform. Sea level rose through the Middle Miocene, began significantly fluctuating until the end of the Pleistocene, and rose in the Holocene to present sea level. Initially, in the Early Miocene, the siliciclastics were deposited interbedded and mixed with carbonates in northern Florida while carbonates continued to dominate in southern Florida (Scott 1988). By the Middle Miocene, with continued sea-level rise, siliciclastics replaced carbonate deposition (Scott 1988; Missimer 2002). Carbonate deposition continued only in the southernmost portions of the platform, and siliciclastic sediments continued to be transported further south and, ultimately, dominated the deposition system on most of the Florida Platform by the early Pliocene. Carbonates continued to

be produced but on a much more limited scale and in the late Neogene, carbonate most often occurred as matrix. Siliciclastic sediments prograded onto the southernmost portion of the platform in the Pliocene, forming the foundation for the northern half of the Florida Keys (Cunningham et  al. 1998). In the Quaternary, siliciclastics dominated over much of the platform. However, in the late Quaternary, with a reduction in siliciclastic supply, carbonate deposition began to occur over portions of the southernmost peninsula. Sediments deposited in the Miocene covered the entire platform; however, subsequent erosion and redeposition created the distributional pattern seen today (Scott et al. 2001). The initial distribution of Pliocene sediments is not known but can reasonably be inferred to have been more extensive than the present occurrence (Scott et al. 2001). Unusual depositional environments are recorded on the Florida Platform in the late Cenozoic (Neogene) as the result of sea-level fluctuations and marine upwelling bottom waters. Major phosphorite and palygorskite deposits formed as the result of these conditions (Weaver and Beck 1977; Riggs 1979; Scott 1988; Compton 1997). The age of the phosphorites indicate that the phosphogenic environment occurred in the Early and Middle Miocene (Compton 1997). The peri-marine environments in which the palygorskite deposits formed also occurred during the Miocene in northwestern Florida (Weaver and Beck 1977). Palygorskite also formed in alkaline lakes in the western-central part of the peninsula in association with sea-level fluctuations (Upchurch et al. 1982). In the late Neogene and into the Quaternary, climate and depositional conditions allowed the development of extremely fossiliferous molluscan-bearing lithologic units. Some of the formations defined within the late Neogene and early Quaternary contain some of the most diverse faunas in the world. How these units formed has been a source of discussion (Allmon 1992; Scott and Allmon 1992). Due to the abundance and diversity of the molluscan fossils, paleontologists have been drawn to study these sediments for more than a century (Scott 1997). As sea level rose in the Pleistocene, sediments were deposited over that portion of the platform that is below 18.3–30.5 m (60–100 ft) above sea level. The Pleistocene sea level rose no higher than this level (Colquhoun et al. 1968). The rising sea level in the late Pleistocene and increased carbonate production on the southern portion of the platform allowed for the development of Miami Limestone (Fig. 2.6), a broad carbonate bank and oolite shoal complex, and Key Largo Limestone, the paleo-reef

26 ~ Scott

Figure 2.8. Generalized geologic map of Florida (modified after Scott et al. 2001).

tract of Florida Keys. The Neogene–Quaternary sediments range in thickness from 0 to more than 914  m (3000 ft) (Miller 1986). During the last glacial stage of the Pleistocene, sea level dropped approximately 122 m (400 ft) exposing vast portions of the Florida Platform that are presently beneath marine waters of the Gulf and Atlantic Ocean. Stream and river channels that can be seen on bathymetric maps provide evidence for erosion during sea-level lowstands. Holocene sea level rose from approximately 18  m (60 ft) depth to the present level, and 8000 to 6000 years BP-archeological sites are found offshore in the Florida Big Bend (Faught and Donoghue 1997). Davis (1997)

stated that the 3000 years BP-sea level was not significantly lower than the present sea level. Davis believes that much of the present-day coastline formed during the last 3000 years as the result of the relatively stable sealevel conditions. The Florida Everglades formed during this general time frame through the deposition of mangrove peat and freshwater calcitic mud covering a broad expanse of Miami Limestone. The distribution of the Neogene and Quaternary units overlying the Paleogene sediments are shown in cross sections in Figure 2.7. A generalized geologic map of Florida is shown in Figure 2.8. The Neogene and Quaternary lithostratigraphic units occurring in the surface and shal-

Geology of the Florida Platform ~ 27

Figure 2.9. Paleogene to Quaternary stratigraphic chart of Florida showing the lithostratigraphic units occurring in the shallow subsurface and at the surface (Scott et al. 2001).

low subsurface of the panhandle, northern, and southern portions of Florida are shown in Figure 2.9.

Hydrogeology The Cenozoic sediments of Florida form a series of aquifer systems, which provide more than 90% of the drinking

water for the state (Berndt et al. 1998). The aquifer systems are the Floridan, intermediate, and surficial (Southeastern Geological Society [SEGS] 1986; see Miller 1986 and Arthur et al. 2008 for overviews). The Floridan aquifer system (FAS) is composed of Paleogene carbonates with highly variable permeability. This aquifer system, which is widespread in the southeastern United States, is one of the most productive

28 ~ Scott

aquifers in the world (Miller 1986; Berndt et al. 1998). Budd and Vacher (2004) characterize the Floridan as a multi-porosity aquifer: a fractured, porous aquifer where confined, and a karstic, fractured, porous aquifer where unconfined. The FAS occurs over the entire platform. The base of the FAS occurs in the lower Paleogene rocks where evaporites restrict the permeability (Miller 1986; SEGS 1986). The top occurs where the carbonates are overlain by impermeable sediments of the intermediate aquifer system or by surface sands. The intermediate aquifer system (IAS) (referred to by the SEGS [1986] as the intermediate aquifer system / confining unit) is composed of permeable and impermeable sediments deposited during the Neogene. The siliciclastics flooding onto the Florida Platform during the Miocene and Pliocene contained an abundance of clay. Deposition of the clayey sediments on the Paleogene carbonates created an impermeable sequence of confining beds (Miller 1986, 1997). Permeable carbonate and siliciclastic sediments are, in some areas, interbedded with the impermeable units creating regionally limited water-producing zones (Miller 1986, 1997). The base of the IAS occurs at the top of the regionally extensive, permeable carbonates of the FAS (SEGS 1986). The top of the IAS is placed at the top of the laterally extensive and vertically persistent lower-permeability beds (SEGS 1986). The IAS is absent over much of the Ocala Platform. The surficial aquifer system (SAS) is composed of late Pliocene through the Pleistocene–Holocene, permeable siliciclastic and carbonate sediments with some zones of more clayey, less-permeable sediments (Berndt et al. 1998). In 2 areas of the state, the SAS is particularly important since the FAS does not contain potable water. In these areas, the westernmost panhandle and southeastern peninsula, the SAS is the primary source of drinking water. In the western panhandle, the SAS is a thick sequence (up to 152 m [500 ft]) of siliciclastic sediments (Sand and Gravel Aquifer). In the southeastern peninsula, the SAS is made of very permeable, interbedded carbonates and siliciclastics, which underlie some of the largest metropolitan areas in Florida (Biscayne Aquifer). The base of the SAS occurs at the top of the laterally extensive and vertically persistent lower-permeability beds (SEGS 1986). The SAS is generally absent on the Ocala Platform.

Geomorphology The Florida Platform extends southward from the continental United States separating the Gulf of Mexico from

the Atlantic Ocean. The exposed portion of the platform, the Florida Peninsula, constitutes approximately onehalf of the Florida Platform measured between the 200-m (600 ft) depth contour of the continental shelves. The axis of the platform extends northwest to southeast approximately along the present-day west coast of the peninsula. The Florida Peninsula, from the St. Mary’s River to Key West, measures nearly 725 km (450 mi). From the Alabama–Florida line to the Atlantic coastline is approximately 595 km (370 mi). Florida lies entirely within the Coastal Plain Physiographic Province as defined by Fenneman (1938) and is the only state in the United States that falls completely within the Coastal Plain. Much of the surface of Florida shows the influence of the marine processes that transported and deposited the later Tertiary, Quaternary, and Holocene sediments. Fluvial processes, although more important in the panhandle, have helped sculpt the entire state, particularly during the lowstands of sea level, redistributing the marine sediments. Karst processes have had a dramatic effect on the Florida landscape due to the near-surface occurrence of soluble carbonate rocks. Middle Eocene to Pleistocene carbonate sediments are affected by karstification over large areas of the state. Siliciclastic sediments, ranging in thickness from a 1 m (3 ft) to more than 61 m (200 ft), overlie the karstified carbonates. More than 700 springs are recognized in Florida with the major springs occurring within the karstic areas of the state (Scott et al. 2004). The vast majority of the springs are located in the Ocala Karst District, the Central Lake District, and the Dougherty Karst Plain District (Scott unpublished). The general geomorphology of Florida consists of east– west trending highlands in the northern and western portions of the state and north–south trending highlands extending approximately two-thirds the length of the peninsula. Coastal lowlands occur between the highlands and the coastline that wraps around the entire state. The highest point in the state, 105 m (345 ft) above sea level, occurs in the Western Highlands near the Alabama–Florida state line in Walton County. There are several hilltops in the Central Highlands that exceed 91 m (300 ft) msl in elevation. Florida has the distinction of having the lowest high point of any state in the United States. White et al. (1964) and White (1958, 1970) delineated the geomorphic subdivisions that most geologists working in the state recognize (see Schmidt 1997 for a review). Scott is creating a new geomorphic map of the state (unpublished).

Geology of the Florida Platform ~ 29

References Allmon, W. D. 1992. Whence southern Florida’s PlioPleistocene shell beds? Pp.1–20 in T. M. Scott and W. D. Allmon (eds.), Plio-Pleistocene Stratigraphy and Paleontology of Southern Florida. Special Publication 36. Tallahassee, Fla.: Florida Geological Survey. Applegate, A. V., G. O. Winston, and J. G. Palacas. 1981. Subdivision and regional stratigraphy of the pre-Punta Gorda rocks (lowermost Cretaceous-Jurassic?) in south Florida. Gulf Coast Association of Geological Societies Transactions 31:447–53. Arthur, J. D. 1988. Petrogenesis of the Early Mesozoic Tholeiite in the Florida Basement and an Overview of Florida Basement Geology. Report of Investigation No. 97. Tallahassee, Fla.: Florida Geological Survey. 39 pp. Arthur, J. D., C. Fischler, C. Kromhout, J. M. Clayton, G. M. Kelley, R. A. Lee, L. Li, M. O’Sullivan, R. C. Green, and C. Werner. 2008. Hydrogeologic Framework of the Southwest Florida Water Management District. Bulletin 68. Tallahassee, Fla.: Florida Geological Survey. 102 pp. Barnett, R. S. 1975. Basement structure of Florida and its tectonic implications. Gulf Coast Association of Geological Societies Transactions 25:122–42. Berndt, M. P., E. T. Oaksford, and G. L. Mahon. 1998. Groundwater. Pp. 38–63 in E. A. Fernald and E. D. Purdum (eds.), Water Resources Atlas of Florida. Tallahassee, Fla.: Florida State University. Bourgeois, J., T. A. Hansen, P. L. Wiberg, and E. G. Kauffman. 1988. A tsunami deposit at the Cretaceous-Tertiary boundary in Texas. Science 241:567–70. Braunstein, J., P. Huddlestun, and R. Biel. 1988. Gulf Coast Region, Correlation of Stratigraphic Units of North America (COSUNA) Project. Correlation Chart. Tulsa, Okla.: American Association of Petroleum Geologists. Brewster-Wingard, G. L., T. M. Scott, L. E. Edwards, S. D. Weedman, and K. R. Simmons. 1997. Reinterpretation of the peninsular Florida Oligocene: an integrated stratigraphic approach. Sedimentary Geology 108:207–28. Bryan, J. R. 1991. Stratigraphic and paleontologic studies of the Paleocene and Oligocene carbonate facies of the eastern Gulf Coastal Plain. Ph.D. diss., University of Tennessee, Knoxville, Tenn. 324 pp. Budd, D. A., and H. L. Vacher. 2004. Matrix permeability of the confined Floridan Aquifer, Florida, USA. Hydrogeology Journal 12:531–49. Chowns, T. M., and C. T. Williams. 1983. Pre-Cretaceous rocks beneath the Georgia Coastal Plain: Regional implications. Pp. L1-L42 in G. S. Gohn (ed.), Studies Related to the Charleston, South Carolina, Earthquake of 1886: Tectonics

and Seismicity. Professional Paper 1313-L. Washington, D.C.: U.S. Geological Survey. Colquhoun, D. J., S. M. Herrick, and H. G. Richards. 1968. A fossil assemblage from the Wicomico Formation in Berkeley County, South Carolina. Geological Society of America Bulletin 79:1211–220. Compton, J. S. 1997. Origin and paleogeographic significance of Florida’s phosphorite deposits. Pp. 195–216 in A. F. Randazzo and D. S. Jones (eds.), The Geology of Florida. Gainesville, Fla.: University Press. Cunningham, K., D. McNeill, L. Guertin, P. Ciesielski, T. Scott, and L. de Verteuil. 1998. New Tertiary stratigraphy for the Florida Keys and southern peninsula of Florida. Geological Society of America Bulletin 110:231–58. Cunningham, K., S. D. Locker, A. C. Hine, D. Bukry, J. A. Barron, and L. A. Guertin. 2003. Interplay of late Cenozoic siliciclastic supply and carbonate response on the southeast Florida Platform. Journal of Sedimentary Research 73: 31–46. Dallmeyer, R. D., M. Caen-Vachette, and M. Villeneuve. 1987. Emplacement age of post-tectonic granites in southern Guinea (West Africa) and the peninsular Florida subsurface: Implications for the origins of southern Appalachian exotic terranes. Geological Society of America Bulletin 99:87–93. Davis, R. A. Jr. 1997. Geology of the Florida Coast. Pp. 155–68 in A. F. Randazzo and D. S. Jones (eds.), The Geology of Florida. Gainesville, Fla.: University Press. Duncan, J. G., W. L. Evans III, and K. L. Taylor. 1994. Geologic Framework of the Lower Floridan Aquifer System, Brevard County, Florida. Bulletin 64. Tallahassee, Fla.: Florida Geological Survey. 90 pp. Faught, M. K., and J. F. Donoghue. 1997. Marine inundated archeological sites and paleofluvial systems: Examples from a karst-controlled continental shelf setting in Apalachee Bay, northeastern Gulf of Mexico. Geoarchaeology 12: 417–58. Fenneman, N. M. 1938. Physiography of Eastern United States. New York: McGraw-Hill Book Company, Inc. 714 pp. Guertin, L. A. 1998. A late Cenozoic mixed carbonate / siliciclastic system, south Florida: lithostratigraphy, chronostratigraphy and sea-level record. Ph.D. diss., University of Miami, Miami, Fla. 424 pp. Guertin, L. A., T. M. Missimer, and D. F. McNeill. 2000. Hiatal duration of correlative sequence boundaries from Oligocene-Pliocene mixed carbonate / siliciclastic sediments of the south Florida Platform. Sedimentary Geology 134:1–26. Heatherington, A. L., and P. A. Mueller. 1997. Geochemistry and origin of Florida crustal basement terranes. Pp. 27–37

30 ~ Scott

in A. F. Randazzo and D. S. Jones (eds.), The Geology of Florida. Gainesville, Fla.: University Press. Hildebrand A. R., G. T. Penfield, D. A. Kring, M. Pilkington, Z. A. Camargo, S. A. Jacobsen, and W. V. Boynton. 1991. Chicxulub crater: a possible Cretaceous / Tertiary boundary impact crater on the Yucatan Peninsula, Mexico. Geology 19:867–71. Huddlestun, P. F. 1993. The Oligocene. Bulletin 105. Atlanta, Ga.: Georgia Geologic Survey. 152 pp. Jones, D. S. 1997. The marine invertebrate fossil record of Florida. Pp. 89–117 in A. F. Randazzo and D. S. Jones (eds.), The Geology of Florida. Gainesville, Fla.: University Press. Kring, D. A. 2000. Impact events and their effect on the origin, evolution and distribution of life. GSA Today 10:1–7. Leve, G. W. 1966. Ground Water in Duval and Nassau Counties, Florida. Report of Investigations No. 43. Tallahassee, Fla.: Florida Geological Survey. 91 pp. Lichtler, W. F., W. Anderson, and B. F. Joyner. 1968. Water resources of Orange County, Florida. Report of Investigations No. 50. Tallahassee, Fla.: Florida Geological Survey. 150 pp. Lloyd, J. M. 1989. 1986 and 1987 Florida Petroleum Production and Exploration. Information Circular 106. Tallahassee, Fla.: Florida Geological Survey. 39 pp. ———. 1997. 1994 and 1995 Florida Petroleum Production and Exploration. Information Circular 111. Tallahassee, Fla.: Florida Geological Survey. 62 pp. Matsui, T., F. Imamura, E. Tajika, Y. Nakano, and Y. Fujisawa. 1999. K-T impact tsunami. Abstract 1527 in Lunar and Planetary Science Confernce 30. Houston, Tex.: Lunar and Planetary Institute. CD-ROM. McBride, J. H., and K. D. Nelson. 1988. Integration of COCORP deep reflection and magnetic anomaly analysis in the southeastern United States: Implications for the origin of the Brunswick and East Coast magnetic anomalies. Geological Society of America Bulletin 100:436–45. McFarlan, E., and Menes, L. S. 1991. Lower Cretaceous. Pp. 181–204 in A. Salvador (ed.), The Geology of North America, Volume J, The Gulf of Mexico Basin. Boulder, Colo.: Geological Society of America. Miller, J. A. 1986. Hydrogeologic Framework of the Floridan Aquifer System in Florida, and in Parts of Georgia, Alabama and South Carolina. Professional Paper 1403-B. Washington, D.C.: U. S. Geological Survey. 91 pp. ———. 1997. Hydrogeology of Florida. Pp. 69–88. in A. F. Randazzo and D. S. Jones (eds.), The Geology of Florida. Gainesville, Fla.: University Press. Missimer, T. M. 2001. Siliciclastic facies belt formation and the Late Oligocene to Middle Miocene partial drowning of the

southern Florida Platform. Gulf Coast Association of Geological Societies Transactions 51:229–38. ———. 2002. Late Oligocene to Pliocene Evolution of the Central Portion of the South Florida Platform: Mixing of Siliciclastic and Carbonate Sediments. Bulletin 65. Tallahassee, Fla.: Florida Geological Survey. 184 pp. Missimer, T. M., and R. G. Maliva. 2004. Tectonically induced fracturing, folding and groundwater flow in south Florida. Gulf Coast Association of Geological Societies Transactions 54:443–59. Missimer, T. M., R. G. Maliva, C. W. Walker, and E. Owosina. 2000. Anatomy of a nearshore mixed siliciclastic-carbonate deposit, the Plio-Pleistocene of southern Broward County, Florida. Gulf Coast Association of Geological Societies Transactions 50:111–28. Opdyke, N. D., D. S. Jones, B. J. MacFadden, D. L. Smith, P. A. Muller, and R. D. Schuster. 1987. Florida as an exotic terrane: Paleomagnetic and geochronologic investigation of the lower Paleozoic rocks from the subsurface of Florida. Geology 15:900–903. Pirkle, W. A. 1970. The offset course of the St. Johns River, Florida. Southeastern Geology 13:39–59. Randazzo, A. F. 1997. The sedimentary platform of Florida: Mesozoic to Cenozoic. Pp. 39–56 in A. F. Randazzo and D. S. Jones (eds.), The Geology of Florida. Gainesville, Fla.: University Press. Riggs, S. R. 1979. Phosphorite sedimentation in Florida—a model phosphogenic system: Economic Geology 74: 285–314. Salvador, A. 1991a. Introduction. Pp. 1–12 in A. Salvador (ed.), The Geology of North America, Volume J, Gulf of Mexico Basin. Boulder, Colo.: Geological Society of America. ———. 1991b. Triassic-Jurassic. Pp. 131–80 in A. Salvador (ed.), The Geology of North America, Volume J, Gulf of Mexico Basin. Boulder, Colo.: Geological Society of America. Schlee, J. S., W. Manspeizer, and S. R. Riggs. 1988. Paleoenvironments: offshore Atlantic U.S. margin. Pp. 365–85 in R. E. Sheridan and J. A. Grow (eds.), The Geology of North America, Volume 1–2, The Atlantic Continental Margin. Boulder, Colo.: Geological Society of America. Schmidt, W. 1984. Neogene Stratigraphy and Geologic History of the Apalachicola Embayment. Florida. Bulletin 58. Tallahassee, Fla.: Florida Geological Survey. 146 pp. ———. 1997. Geomorphology and physiography of Florida. Pp. 1–12. in A. F. Randazzo and D. S. Jones (eds.), The Geology of Florida. Gainesville, Fla.: University Press. Schultz, P. H. 1996. Cretaceous-Tertiary (Chicxulub) impact angle and its consequences. Geology 24:963–67.

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Scott, T. M. 1988. The Lithostratigraphy of the Hawthorn Group (Miocene) of Florida. Bulletin 59. Tallahassee, Fla.: Florida Geological Survey. 148 pp. ———. 1992. A Geological Overview of Florida. Open File Report 50. Tallahassee, Fla: Florida Geological Survey. 78 pp. ———. 1997. Miocene to Holocene history of Florida. Pp. 57–67 in A. F. Randazzo and D. S. Jones (eds.), The Geology of Florida. Gainesville, Fla.: University Press. ———. 2001. Text to Accompany the Geologic Map of Florida. Open File Report 80. Tallahassee, Fla.: Florida Geological Survey. 29 pp. Scott, T. M., and W. D. Allmon. 1992. Plio-Pleistocene stratigraphy and paleontology of southern Florida. Pp. 21–26 in T. M. Scott and W. D. Allmon (eds.), Plio-Pleistocene Stratigraphy and Paleontology of Southern Florida. Special Publication 36. Tallahassee, Fla.: Florida Geological Survey. Scott, T. M., G. H. Means, R. P. Meegan, R. C. Means, S. B. Upchurch, R. E. Copeland, J. Jones, T. Roberts, and A. Willett. 2004. Springs of Florida. Bulletin 66. Tallahassee, Fla.: Florida Geological Survey. 377 pp. Scott, T. M., K. M. Campbell, F. R. Rupert, J. D. Arthur, R. C. Green, G. H. Means, T. M. Missimer, J. M. Lloyd, J. W. Yon, and J. D. Duncan. 2001. Geologic Map of the State of Florida. Map Series 146. Tallahassee, Fla.: Florida Geological Survey. Smith, D. L. 1982. Review of the tectonic history of the Florida basement: Tectonophysics. 88:1–22. Smith, D. L., and K. M. Lord. 1997. Tectonic evolution and geophysics of the Florida basement. Pp. 13–26 in A. F. Randazzo and D. S. Jones (eds.), The Geology of Florida. Gainesville, Fla.: University Press. Sohl, N. F., E. Martinez, P. Salmeron-Urena, and F. SotoJaramillo. 1991. Upper Cretaceous. Pp. 205–44 in A. Salvador (ed.), The Geology of North America, Volume J, Gulf of Mexico Basin. Boulder, Colo.: Geological Society of America. Southeastern Geological Society. 1986. Hydrogeological Units of Florida. Special Publication 28. Tallahassee, Fla.: Florida Geological Survey. 8 pp. Stuckey, J. L. 1965. North Carolina: Its Geology and Mineral Resources. Raleigh, N.C.: North Carolina Department of Conservation and Natural Resources. 550 pp. Thomas, W. A., T. M. Chowns, D. L. Daniels, T. L. Neathery,

L. Glover III, and R. J. Gleason. 1989. The subsurface Appalachians beneath the Atlantic and Gulf Coastal Plains. Pp. 445–58 in R. D. Hatcher, W. A. Thomas, and G. W. Viele (eds.), The Appalachian-Ouachita Orogen in the United States, The Geology of North America, Volume F-2. Boulder, Colo.: The Geological Society of America. Upchurch, S. B., R. N. Strom, and M. G. Nuckels. 1982. Silicification of Miocene rocks from central Florida. Pp. 251–284 in T. M. Scott, and S. B. Upchurch (eds.), Miocene of the Southeastern United States. Special Publication 25. Tallahassee, Fla.: Florida Bureau of Geology. Warzeski, E. R., K. J. Cunningham, R. N. Ginsburg, J. B. Anderson, and Z. D. Ding. 1996. Neogene mixed siliciclastic and carbonate foundation for the Quaternary carbonate shelf, Florida Keys. Journal of Sedimentary Research 66:788–800. Weaver, C. E., and K. C. Beck. 1977. Miocene of the southeastern United States: A model for chemical sedimentation in a peri-marine environment. Sedimentary Geology 17:1–234. White, W. A. 1958. Some Geomorphic Features of Central Peninsular Florida. Bulletin 41. Tallahassee, Fla.: Florida Geological Survey. 92 pp. ———. 1970. The Geomorphology of the Florida Peninsula. Bulletin 51. Tallahassee, Fla.: Florida Geological Survey. 164 pp. White, W. A., H. S. Puri, and R. O. Vernon. 1964. Physiographic setting. Pp. 7-12 in H. S. Puri and R. O. Vernon (eds.), Summary of the Geology of Florida and a Guidebook to the Classic Exposures. Special Publication 5 (revised). Tallahassee, Fla.: Florida Geological Survey. Winston, G. O. 1987. Generalized stratigraphy and geologic history of the South Florida Basin. Pp. 230–33 in F. J-M R. Maurrasse (ed.), Symposium on South Florida Geology. Memoir 3. Miami, Fla.: Miami Geological Society. ———. 1991. Atlas of Structural Evolution and Facies Development on the Florida-Bahama Platform—Triassic through Paleocene. Miami, Fla.: Miami Geological Society. 39 pp. Woods, R. D., A. Salvador, and A. E. Miles. 1991. Pre-Triassic. Pp. 109–29 in A. Salvador (ed.), The Gulf of Mexico Basin: The Geology of North America, Volume J. Boulder, Colo.: Geological Society of America. Wyrick, G. G. 1960. The Ground Water Resources of Volusia County, Florida. Report of Investigations 22. Tallahassee, Fla.: Florida Geological Survey. 65 pp.

␥3

Pre-Holocene Geological Evolution of the Northern Gulf of Mexico Basin William E. Galloway

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. The Gulf of Mexico basin is a world-class repository of hydrocarbons (Nehring 1991). It has been actively explored for nearly 100 years, creating a three-dimensional well and reflection seismic database of unique abundance, extent, and diversity. 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. This chapter will focus on the history of this northern fill, with emphasis on the area beneath the present continental shelf of the northern Gulf.

Crustal Structure and Basin Origin The Gulf of Mexico basin was created by crustal extension and seafloor spreading during the Mesozoic breakup of Pangea (Sawyer et al. 1991). Most of the structural basin is underlain by transitional crust that consists of continental crust that was stretched and attenuated primarily by Middle to Late Jurassic rifting (Fig. 3.1). The northernmost basin margin is underlain by thick transitional

crust, which displays modest thinning and typically lies at subsea depths between 2 and 10 km. Much of the present lower coastal plain, shelf, and continental slope is underlain by thin transitional crust, which is generally less than half of the 35-km thickness typical of continental crust and is buried more than 10 km below sea level. Basement may lie below 20 km in the depocenter beneath the southern Louisiana coastal plain and adjacent continental shelf (Peel et al. 1995). 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), if not agreed upon in detail. The Gulf of Mexico opened by the separation of the North and South American plates. Triassic through early Jurassic tensional deformation created a series of basement grabens and half grabens, which are filled with terrestrial red beds and volcanics. Crustal stretching in Bathonian and Callovian time initiated a broad sag, which opened initially 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. A regional unconformity beneath the evaporite layer separates localized syn-rift from blanket post-rift deposits (Sawyer et al. 1991; Buffler and Thomas 1994). Opening of the Gulf entailed approximately 500 km of extension accompanied by southward migration and counterclockwise rotation of the Yucatan block 33

34 ~ Galloway

95o Ouachita Mountains

90o

SA

2

ETB

MU

NLSB

Pre - marine evaporite Crustal types

1

2

1

85o

Appalachian Mountains

6 10

8 6

MSB

THICK TRANSITIONAL CRUST

12

30 o

12

10

4

14

10

RGE

L

IT

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TR

N HI

6

8

S AN

8

MGA

6

14

TE

14

IO

12 10

S RU

C

NA

6

2

16

T

8

1

AE

14

SMA

10 10

12

SrA

12

T

4

25 o

2 WA

1

Llano Uplift

8

10

C

8

4

US

14

CR

AL

8

N

6

NI

C

S RU

IO

BA

T SI

EA

MEXICO

T

12

T

TA

CU

4

OC

AN

CK

TR

2

10

I

12

TH

YUCATAN

0

100 km

Figure 3.1. Crustal types, depth to basement (in meters), and original distribution of Jurassic Louann salt (pre-marine evaporite) beneath the Gulf of Mexico basin. Principal basement structural elements 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 (modified from Sawyer et al. 1991).

and by extensive north northwest–south southeast shear along the east and west flanks of the basin (Marton and Buffler 1993, 1999; Pindell and Kennan 2001). Crustal rupture began by the Oxfordian and continued until the termination of spreading in the latest Berriasian or early Valanginian. By the end of the Early Cretaceous, combined deposition and subsidence had created the modern morphology of the Gulf basin (Winker and Buffler 1988). Late Cretaceous and, especially, Cenozoic history was dominated by loading subsidence.

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 juxtaposes thick transitional and oceanic crust. This pronounced asymmetry suggests a simple-shear model for extension (Marton and Buffler 1993; Watkins et al. 1995). Additional Mesozoic and Cenozoic tectonic phases have influenced subregional subsidence history of the Gulf. Several of the marginal highs, including the San Marcos arch, Sabine arch, and Monroe uplift, display

Geological Evolution ~ 35

short pulses of uplift of as much as several hundred meters (Laubach and Jackson 1990). Feng et  al. (1994) documented accelerated early Cenozoic tectonic subsidence in the western deep Gulf and suggested a Laramide foreland basin overprint.

Structural Framework The depositional history of the Gulf of Mexico basin is affected both by basement structure, which subtly influenced sediment supply and accumulation patterns, and gravity tectonics. Dynamic interactions among depositional loading and sediment and salt mobilization rearranged distribution of accommodation space and deformed buried strata.

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. 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 northwest–southeast trending basement transfer faults created during Atlantic and Gulf extension and spreading phases (Watkins et al. 1995; Huh et al. 1996). These lineaments influenced deposition of salt, which, in turn, has arguably controlled subsequent location and pattern of salt structures and subsequent generation and migration of hydrocarbons beneath the modern shelf and slope (Stephens 2001).

Gravity Tectonic Structures The Gulf of Mexico basin fill displays one of the best described and most complex assemblages of gravity tectonic structures found in the world (Winker 1982; Jackson et  al. 1994; Diegel et  al. 1995; Jackson 1995; Peel et al. 1995; Schuster 1995; Watkins et al. 1996a; Watkins et al. 1996b; Rowan et al. 1999; 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 basinal mud up section and basinward throughout Gulf history. The panoply of structures includes:

1. Growth fault families and related structures, including the primary synthetic growth fault, splay faults, antithetic faults, and rollover anticlines. 2. Allochthonous salt bodies, including salt canopies and salt sheets. 3. Salt welds that juxtapose discordant stratigraphies where nearly complete expulsion of salt stock feeder dikes, salt tongues, or salt canopies has occurred. 4. Roho fault families where lateral salt-tongue extension by gravity spreading has created a linked assemblage of extensional faults and compensating, downslope compressional toe faults, anticlines, and salt injections in the overlying sedimentary cover. 5. Salt diapirs and their related withdrawal synclines and minibasins where depositional loading of the autochthonous Louann or allochthonous salt canopies created high-relief diapirs and intervening sediment-filled depressions. 6. Compressional fold belts (Weimer and Buffler 1992; Fiduk et al. 1995; Trudgill et al. 1999) typically formed at the base of the slope and on the basin plain where a stepped discontinuity or termination of the decollement layer occurs. In addition, the curvature of the northwestern Gulf margin created northwest–southeast compression that is manifested in Miocene strata beneath the Texas continental slope.

Growth Structure Domains A complex array of gravity tectonic structures lies within the Cenozoic sedimentary wedge of the northern Gulf of Mexico basin (Fig. 3.2). Principal structural domains include an inboard series of strike-aligned shale and saltdetached growth fault families beneath the coastal plain and Texas shelf, roho fault families beneath southern 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 buried basin-floor compressional fold belts. Sediment loading of the salt canopies has created a series of minibasins that are largely filled beneath the outer shelf but that persist as closed bathymetric lows on the continental slope. Each structural domain had a defined growth phase associated with successive episodes of Cenozoic sediment accumulation in the Gulf. Domains generally

36 ~ Galloway

100°

95° Sta

90° Pic kin s

te Line FZ

ET SDB

Zo n - T e alc o

Fault

Zon e

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lt au sF e n M lco Ba g li n u L

ilb er t on F.Z.

Edg e

SM

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re t

ex ia

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us eo ac

-G

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Shelf

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OMC

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SC Fig. 4A Fig. 4B Fig. 3 Bounding Graben Faults

Salt Dome Basins SDB

Wilcox Detachment

Upper Eocene Detachment

WD

UED

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Mixed Upper Eocene and Top Salt Detachment

Vicksburg Detachment

Oligocene - Miocene Detachment

Shelf Minibasins

MD

VD

OMD

SM

Roho

Salt Canopy

Oligocene-Lower Miocene Compression

Miocene Compression

R

SC

OMC

MC

Figure 3.2. Structural domains of the northern Gulf of Mexico. Compiled from Ewing (1991), Diegel et al. (1995), and Jamieson et al. (1998). DCSDB, De Soto Canyon salt basin; MSDB, Mississippi salt dome basin; NLSDB, North Louisiana salt dome basin; ETSDB, East Texas salt dome basin; FZ, fault zone; LK, Lower Cretaceous; UK, Upper Cretaceous.

become younger basinward, beginning with the Paleocene Wilcox 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. The three-dimensional structural and stratigraphic architectures of the northern basin are illustrated by a regional north-south section across the north-central basin fill (Fig. 3.2). The boundary between thick and thin transitional crust created a subsidence hinge that focused development of the Cretaceous continental shelf margin as an extensive reef system. Basinward, the thick Cenozoic sedimentary prism overlies thin transitional crust, which has been depressed as much as 20 km by sedimentary loading. The prism extends beneath the modern coastal plain and shelf, reaching its thickest point near the present continental margin. The modern continental slope extends basinward to about the position of the transitional-oceanic crust boundary. Beneath this sedi-

ment 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. Off-stepping deposition acted as a giant rolling pin, pushing salt upward and basinward into 3 major salt canopies. The subshelf canopies were loaded and largely evacuated by subsequent deposition, forming the vast central Gulf shelf minibasin and roho domains (Fig. 3.2). Beneath the continental slope, a shallow salt canopy forms the primary slope minibasin and salt canopy domains, which terminate in the Sigsbee Scarp. 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 (Fig. 3.3). Transects through the northeastern and northwestern Gulf margins (Fig. 3.4) illustrate features of additional

Geological Evolution ~ 37

N

Kr reef

Paleocene - Miocene Growth Faults

S

Slope Minibasins

Shoreline

km 0

Sigsbee Scarp SL

5 10

TTC

15

D

D C

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C

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Oceanic Crust 0

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25 Salt

C

C D

D

Jurassic

U. & L. Cretaceous

Paleo-Eocene

Oligocene

Miocene

100 km

Pliocene

Pleistocene

Figure 3.3. Dip cross section through the north central Gulf of Mexico showing crustal types, generalized stratigraphy, and structural elements. For location see Figure 3.2. Kr, lower Cretaceous; TTC, thick transitional continental crust; SL, sea-level datum (modified from Peel et al. 1995). N

S

Kr reef

km 0

5

10

15

?

Thick Transitional Crust

Thin Transitional Crust

?

0

v.e. ~ 5:1

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Miocene G.F.

km 0

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Port Isabel F.B.

Perdido F.B.

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Salt

Jurassic

Cretaceous

EoceneOligocene

Miocene

v.e. ~ 5:1

50 km

PlioPleistocene

Figure 3.4. Dip cross sections of the northeastern (A) and northwestern (B) Gulf continental margins (modified from Peel et al. 1995). For locations, see Figure 3.2. Kr, lower Cretaceous; G.F., growth fault belt; F.B., fold belt; v.e. = vertical exaggeration.

structural domains and variants on the central basin stratigraphy. In the northeastern Gulf (Fig. 3.4A), the total basin fill is relatively thin, depressing the crust only to depths between 7 and 11 km. The crustal boundary, which here lies beneath the modern shelf, again pins the location of the Cretaceous shelf margin. Neogene deposition prograded the shelf edge about 50 km farther basinward. Growth faults are few. Limited salt stocks rise from the largely evacuated autochthonous Louann. The slope toe intersects the east end of the Miocene compression

domain. The northwestern Gulf transect (Fig. 3.4B) is located beneath the continental shelf and slope. The Miocene growth faults of the Oligocene–Miocene detachment province are rooted in a decollement within deep basinal muds of indeterminate age. Basal Louann salt has been expelled both upward as isolated stocks and basinward to the continental slope toe and beyond, leaving a primary weld. The northwestern Gulf displays complex middle Cenozoic compressional domains beneath the modern shelf and slope, including the Perdido and Port Isabel fold

38 ~ Galloway

belts. The Port Isabel fold belt is linked by a decollement to the Oligocene–Miocene detachment province. The Perdido fold belt lies at the depositional limit of Louann salt (Fiduk et al. 1995).

Depositional Framework The stratigraphic architecture of the northern Gulf of Mexico basin displays many elements typical of divergent continental margins (Winker 1984; Winker and Buffler 1988): (1) Above a breakup unconformity, initial strata onlapped the subsiding basin margin. (2) Following this onlap phase, sediment supply overcame subsidence, and offlap stratigraphy dominates. With offlap, a deep, sediment-starved basin center became separated from the marginal coastal plain and shelf by a clearly defined shelf edge and slope. (3) Episodes of Cenozoic clastic sediment supply deposited a succession of aggrading and offlapping stratal units separated by major marine flooding horizons, sediment-starvation surfaces, and erosional unconformities that completely filled in the proximal basin and overflowed onto the basin floor. The Oligocene–Pleistocene episodes constructed most of the modern continental shelf and slope.

Depositional Episodes and Sequences Using the Frazierian depositional episode 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, where prominent transgressive markers record glacio-eustatic sea-level rise events (Galloway et al. 2000). 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., Mitchum et  al. 1993; Yurewicz et  al. 1993; Mancini and Puckett 1995; Lawless et al. 1997), especially to late Neogene strata

that are strongly influenced by glacio-eustasy (Martin and Fletcher 1995; Weimer et al. 1998). Building upon the syntheses of Winker and Buffler (1988), Galloway (1989b), and Galloway et al. (2000), I propose a simplified genetic stratigraphic framework constructed by 28 composite Gulf of Mexico depositional episodes (Figs. 3.5–3.8). Each episode records a long-term (about 2–12 million years) 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 (Figs. 3.6, 3.8). Deposits of many of the episodes are also bounded by prominent, widely recognized stratigraphic surfaces (Figs. 3.6, 3.8).

Depositional History and Paleogeography The composite 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 (Handford and Loucks 1993; Galloway and Hobday 1996). These systems record geologically long-lived paleogeographic elements of the northern Gulf of Mexico.

Mesozoic Depositional Episodes The upper Jurassic and lowest Cretaceous Louann, Norphlet, Smackover, and Cotton Valley episodes form a tectonostratigraphic megasequence bounded below by the breakup unconformity and above by a prominent intra-Valanginian unconformity, which records the termination of seafloor spreading (Winker and Buffler 1988; Wu et al. 1990; Salvador 1991b; Dodson and Buffler 1997; Marton and Buffler 1999). Following the post-breakup flooding of the shallow Gulf basin through a connection to the Pacific Ocean across central Mexico, widespread deposition of Louann salt blanketed transitional crust (Salvador 1991a, 1991b; Dobson and Buffler 1997). As much as 4 km of nearly pure halite buried the under-

Geological Evolution ~ 39

Time (Ma)

Stages Maastrichtian

Depositional Architecture Escondido

Navarro

Olmos - Nac.

70

Taylor Ariacacho

90

Santonian Coniacian

Austin / Eutaw

Turonian

Eagle Ford

Cenomanian

Buda Kiamichi

Washita Georgetown

CRETACEOUS

110

Woodbine / Tuscaloosa

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

Buckner

Gilmer Smackover

Oxfordian Callovian Middle

160

Late

150

Formation of Oceanic Crust

Bossier

Basalt RAMP

Berriasian

Louann Salt

Bathonian Bajocian

Connection opened to Western Interior Seaway

Figure 3.5. Generalized Mesozoic stratigraphic succession and architecture of the northern Gulf of Mexico basin (modified from Winker and Buffler 1988). Timescale of Gradstein et al. (1995). F.L, Ferry Lake Anhydrite; G.R., Glen Rose Reef.

lying topography and onlapped northward onto the structural margin of the Gulf (Fig. 3.1). Salt accumulation was replaced in the Oxfordian by deposition of a relatively thin, widespread siliciclastic-dominated Norphlet Formation. Eolian, sabkha, and playa deposits indicate continued aridity. Basinward, siliciclastics grade into marine shale and limestone. Oxfordian onlap initiated the first carbonate-

dominated depositional episode. Together, the Smackover, Buckner, and Gilmer formations (Fig. 3.5) record an approximately 5 million year cycle bounded above and below by marine flooding surfaces (Salvador 1991b; Prather 1992; Dobson and Buffler 1997; Goldhammer and Johnson 2000). Initial deposits consisted of fine-grained, dark, carbonate-ramp sediments, which were succeeded by a heterogeneous assemblage of carbonates, including

40 ~ Galloway

Time (Ma)

Stages

Depositional Episodes Clastic Supply Carbonate Platform

Major Surfaces D

Maastrichtian Nacatoch Olmos

70

Composite Episodes Navarro

MFS

San Miguel Campanian UPPER

80

D

Taylor

D Santonian Coniacian

D

Eutaw

90 Turonian

MFS

Austin

D D

Tusc.Woodbine

Tuscaloosa Woodbine

Cenomanian

110

Albian

Washita D

L. Stuart City

Fredericksburg

Paluxy

D Glen Rose

LOWER

Glen Rose

MFS

James 120

D

U. Stuart City

CRETACEOUS

100

D

MFS

Aptian

James

D

Sligo

Sligo

Barremian Hosston

MFS

Hauterivian

130

D Lower Hosston

Valanginian 140

JURASSIC

D Cotton Valley

Cotton Valley D

Kimmeridgian Oxfordian

Haynesville

MFS Smackover

Norphlet

Callovian LOWER

170

UPPER

Tithonian 150

160

Knowles

Berriasian

Smackover MFS

Norphlet Louann

Bathonian

Bojocian

Figure 3.6. Mesozoic depositional episodes of carbonate and siliciclastic sediment accumulation in the northern Gulf of Mexico. Major stratigraphic surfaces include regional erosional unconformities (shown by wavy lines) and maximum flooding disconformities (MFS and arrows). Composite episodes created regionally concordant stratigraphic units bounded by the major surfaces and record long periods of relatively stable paleogeographies. Locally, angular unconformities (shown by D and sloping half arrow) also separate many episode strata. Timescale is from Gradstein et al. (1995). MFS, maximum flooding surface; D, deepening interval.

prominent ramp-edge grain shoals. Seaward, carbonate muds formed a broad carbonate ramp, or eastward a nascent carbonate slope. Clastic influx (Haynesville Formation) was minor. Sandstones of the Cotton Valley depositional episode

(Figs. 3.5, 3.6) abruptly override transgressive Gilmer and Haynesville strata (Salvador 1991b; Dobson and Buffler 1997; Goldhammer and Johnson 2000). Large, sandy delta systems prograded into the eastern Texas basin, Mississippi salt basin, and Apalachicola embayment.

Geological Evolution ~ 41

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

Sparta Queen City

PERCHED

Eocene

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 3.7. Generalized Cenozoic stratigraphic succession and architecture of the northern Gulf of Mexico basin. Timescale of Berggren et al. (1995). E, early; M, middle; L, late; RMP, ramp type margin.

Suspended sediment spread basinward to form a broad, muddy, marine shelf platform. The Cotton Valley depositional episode ended with transgression (Knowles limestone), and its deposits are separated from strata of the Lower Cretaceous Hosston episode by a prominent subaerial unconformity throughout the northern Gulf margin (Salvador 1991b; Goldhammer and Johnson 2000). Updip, this unconformity records the entire Valanginian. Basinward, Valanginian strata form a fore-shelf lowstand wedge (Fig. 3.5). Temporal coincidence of the unconformity with interpreted termination of seafloor spreading in the Gulf suggests that it records intraplate stress regime reorganization that accompanied death of the spreading center. Following the end of Gulf spreading, strata of 6 composite depositional episodes (Fig. 3.6) record diminishing continental source 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; Badalí 2002; Mancini and Puckett 2002). The climatic setting remained tropical and arid. Clastic input decreased and carbonate deposi-

tion dominated the northern Gulf of Mexico (Fig. 3.6). Two phases of regional progradation of the reef-rimmed carbonate margin, separated by a regional early Albian flooding event (Fig. 3.5), produced a well-defined Sligo and Stuart City shelf edge separating open-to-restricted, shallow-platform depositional systems from steep slope and southward-deepening basinal equivalents. Beginning in the middle Albian, the Gulf of Mexico was connected to the Western Interior seaway. The Washita depositional episode was characterized by climax, aggradational growth of the Stuart City reef. The episode terminated with the Mid-Cretaceous unconformity (MCU) (Buffler 1991). This intra-Cenomanian unconformity marks a basin-scale reorganization of regional depositional patterns. The shallow-water carbonate factory was progressively drowned by the shelf deepening, 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). Continental uplift and erosion that supplied clastics have been associated with subcrustal passage of the Ber-

42 ~ 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

Jackson

L

30

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 3.8. Cenozoic depositional episodes of siliciclastic sediment accumulation in the northern Gulf of Mexico. Major stratigraphic surfaces include regional erosional unconformities and maximum flooding disconformities. Composite episodes created regionally concordant stratigraphic units bounded by the major surfaces and record long periods of relatively stable paleogeographies. Neogene episodes incorporate multiple glacio-eustatic cycles. E, early; M, middle; L, late; MFS, maximum flooding surface; Bul. 1, Buliminella 1.

muda hotspot (Cox and Van Arsdale 2002). The resultant Tuscaloosa–Woodbine sequence contains major deltaic systems in the Mississippi embayment and East Texas salt basin. As regressive clastic systems prograded over the dying Stuart City reef, sedimentary bypass and slumping rejuvenated clastic input to the deep northeastern Gulf. Following late Turonian submergence of the northern Gulf margin, Late Cretaceous strata blanketed the foundered shelf edge, subduing its morphology and creating a ramp-like shelf margin. Extensive deep carbonate shelves extended to and beyond the present outcrop. The north-

western Gulf remained an open platform initially connected to the Cretaceous Interior seaway. Up to several hundred meters of Austin strata are found on the northern shelf, but the deep, central Gulf was largely sediment starved during this interval of regional highstand. Upper Taylor episode deposition (Fig. 3.6) records a modest return of terrigenous sediment to the Gulf margin (Sohl et al. 1991) and closure of the connection to the Western Interior seaway. Several unconformities within and at the base and top of the Navarro Group record influence of Laramide crustal stresses on local uplift and sub-

Geological Evolution ~ 43

sidence across the northern Gulf basin (Sohl et al. 1991). Throughout the late Cretaceous the area beneath the modern shelf and slope was dominated by deposition of basinal marls and mudrocks. The Cretaceous–Tertiary boundary strata of the Gulf of Mexico constitute a condensed horizon and also record a cataclysm of global proportions, the Chicxulub meteorite impact event (Hildebrand et al. 1991).

Cenozoic Depositional Episodes The Cenozoic depositional history of the northern Gulf basin has been synthesized by Galloway et al. (2000). Here, I have simplified their 18 northern Gulf of Mexico depositional episodes to 13 (Fig. 3.7) by combining some minor episodes and emphasizing only the first-order changes in supply history and paleogeography. These composite episodes can be further grouped into 5 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 episodes that reflect rejuvenation of western interior drainage basins overprinted by cyclic glacio-eustatic sealevel change. (5) Late Pliocene–Quaternary episodes that record climate deterioration, continental glaciation, and high-amplitude, high-frequency, glacio-eustatic sea-level change.

Continental-scale bedload-dominated fluvial systems spilled first across the East Texas and Louisiana coastal plains and then across the middle Texas coastal plain. Rapid sediment loading mobilized the underlying overpressured muds and Louann salt, initiating growth faults of the Wilcox detachment province (Fig. 3.2) along the paleo-shelf margin. Loading also triggered 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 formed. Ongoing seismicity associated with foreland deformation of the western Gulf margin triggered frequent slumps and slides along the prograding clastic shelf margin from southern Texas to central Louisiana. Several of these slumps nucleated submarine canyons that excavated hundreds of meters of older Wilcox strata (Galloway et al. 1991). Regional transgressions, known by the Yoakum shale and Big shale markers, that punctuated Wilcox deposition are associated with the largest submarine canyons. Progradation into the slumps and canyon heads supplied basin-scale submarine fan and apron systems on the Texas and Louisiana paleo-continental slope. The large, efficient fans spread far out onto the basin floor, beneath the modern shelf and 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 (Fig. 3.7).

Laramide-Related Depositional Episodes Regional flooding of the Gulf margin persisted into the Early Paleocene. Widespread shelf mudrocks and marls of the Midway and Porters Creek formations blanket the northern Gulf. Beginning in the Late Paleocene, 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. Coarse clastic input to the deep basin beneath the modern shelf resumed. Four depositional episodes punctuated Laramiderelated Paleocene through Middle Eocene history of the Gulf of Mexico (Fig. 3.8). The Late Paleocene and Early Eocene Wilcox episodes significantly prograded the northern Gulf continental shelf and slope (Fig. 3.9).

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 the appearance of volcanic ash beds. Yegua and Cockfield fluvial-dominated deltas prograded across the shallow shelf onto the mud-blanketed, distally steepened ramp created by foundering of the relict Upper Wilcox continental margin. The combination of rapid sediment influx and renewed loading triggered a succession of submarine slumps and growth faults along the Yegua delta front (Edwards 1991; Ewing and Vincent 1997).

44 ~ Galloway

100°

95°

90°

85°

JS

LK JS

KWF UK

30°

F

KW

F

QC UW Y

LW

Plio

LM M MU M

F

KS

1

UW

JS

J

UM PB1

Pleist F KW

LK

PB

MM

KW

F

F QC

LK

UM

Maverick Basin (KWF)

LW

UW

el

ann

e Ch

ane Suw

0 0

200 mi 200 km

Large delta systems and subjacent delta - fed aprons

MM

J

Figure 3.9. Compiled paleo-shelf margins of the northern Gulf basin. Margins are shown at their position late in the history of each episode, typically the time of maximum progradation. Location of large delta systems found updip of each clastic margin is shaded. Mesozoic margins: 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. Compiled from Winker and Buffler (1988), Galloway et al. (2000), and unpublished maps.

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 (Gray et al. 2001). Uplift and volcanism directly affected the Gulf by rejuvenation of several drainage basins and by pervasive deposition of easily reworked air-fall ash. The outpouring of recycled sedimentary rocks, volcaniclastics, and reworked, devitrified ash peaked by the mid-Oligocene and continued into the early Miocene. The response in the Gulf was the sediment-supply dominated Frio–Vicksburg depositional episode, which lasted for more than 8 million years (Fig. 3.8). Four delta systems were active during Frio deposition (Galloway et al. 2000). The Oligocene depocenter lies in the Rio Grande embayment of South Texas 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. Norias deposition began with rapid progradation of Early Oligocene Vicksburg deltas onto a 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 con-

sequence was development of the Vicksburg detachment (Fig. 3.2) (Diegel et al. 1995). Subsequent Frio progradation built the continental margin 90 to 145 km basinward (Fig. 3.9), extending the constructional Oligocene slope apron beneath the modern Texas–western Louisiana shelf. Loading of the paleo-continental margin initiated basinward advancing lines of growth faults that form the updip part of the Oligo-Miocene detachment province (Figs. 3.2, 3.3). Beneath the modern south Texas shelf, extension was compensated by faulting and folding along the Port Isabel fold belt, which lay at the base of the Oligocene continental slope, and the Perdido fold belt, located at the basinward termination of salt beneath the abyssal plain (Figs. 3.2, 3.4). The Houston delta is centered beneath the southeastern Texas coastal plain. 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 third delta system, fed by a large, suspended-load fluvial system, prograded across central Louisiana. Between the delta systems, the Frio sequence contains comparably thick successions of strandplain and barrier–lagoon com-

Geological Evolution ~ 45

plexes nourished by longshore reworking of sediment from the deltaic headlands. Between the southeastern Texas and Louisiana delta systems, rapid mid-Oligocene salt withdrawal triggered a brief phase of tilting, collapse, and submarine erosion that interrupted margin progradation (Cossey and Jacobs 1992; Galloway 1998), forming the Hackberry embayment. The northeastern Gulf margin and basin floor remained clastic sediment starved. Decreasing rate of sediment supply in the late Oligocene (Galloway and Williams 1991) terminated the Frio depositional episode, culminating in deposition of the Anahuac shale across the breadth of the northern Gulf margin.

Miocene Depositional Episodes Gulf Miocene stratigraphy is characterized by extensive continental margin progradation (Fig. 3.9). Miocene basin fill reflects 3 (approximately coincident with the Early, Middle, and Late Miocene) multimillion-year depositional episodes (Fig. 3.8) that record the progressive shift of the locus of deposition in the Gulf of Mexico from the northwestern to the northeastern margin. This shift reflects rejuvenation of the Appalachian and Cumberland Plateau uplands. Also during the Miocene, basin-and-range tectonism extended to the western fringe of the Gulf with activation of the Balcones fault system and uplift to the Edwards Plateau. Concurrently, global climate was evolving toward the ice-house world of the late Cenozoic. Increasing amplitude and frequency of glacio-eustatic sea-level fluctuations impacted stratigraphic and facies architecture. The Lower Miocene succession consists of deposits of an 8-million-year episode that followed the Anahuac transgression (Galloway et al. 1986). The Amphistegina shale records the long-term transgression that terminated the composite episode. Following the Anahuac transgression, the bedloaddominated northwestern Gulf fluvial axes continued to wane, although they remained important. Wave reworking and longshore transport dominated the delta systems, shifting the sand depocenter northeast to the laterally adjacent central Texas barrier–strandplain system beneath the inner shelf. Depositional loading caused renewed faulting along the Oligocene–Miocene detachment province (Fig. 3.2). Compression continued along both the Port Isabel and Perdido fold belts. On the central Gulf margin, the paleo-Mississippi and Red rivers continued to grow in importance. At the onset of deposition,

the paleo-continental margin beneath southwestern Louisiana experienced Hackberry-like subsidence, collapse, and mass wasting. Numerous slump scars, fault-expanded shelf-margin deltas, and submarine canyon fills form the resultant “Planulina embayment,” which was nucleated by large-scale salt withdrawal from beneath coastal Louisiana in response to renewed loading. Combined deflation of the shallow Oligocene canopy and extension along the Oligocene and Louann detachment zones (Diegel et al. 1995; Peel et al. 1995) accommodated nearly 7 km of lower Miocene sediment in the central Gulf depocenter. Salt extruded from beneath the prograding margin spread southward, nucleating new canopy complexes beneath the modern Louisiana outer shelf and slope. 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, a delta-fed, muddy shelf encroached eastward, choking reef growth and restricting carbonate platform deposition to the Florida shelf. The middle Miocene sequence records a relatively brief (about 3 million year) 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 (Galloway et  al. 2000; Combellas-Bigott and Galloway 2002a). 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 Tennessee rivers prograded the continental margin as much as 70 km (Fig. 3.9). Initial margin progradation was interrupted, however, by a third pulse of salt evacuation, located beneath the southeastern Louisiana coastal plain (Combellas-Bigott and Galloway 2002a). Beneath the central Texas shelf, the Corsair fluvial-deltaic system prograded the paleo-continental slope beneath the modern central Texas shelf (Galloway et al. 2000). Here, salt withdrawal and prolonged growth of the Corsair fault zone (Oligocene–Miocene detachment province, Fig. 3.2) created a depocenter that was filled by wave-dominated delta and delta-fed apron deposits. 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 northeastern Gulf, combined margin collapse,

46 ~ Galloway

slope bypass, and alignment of a series of dip-elongate slope minibasins created focused submarine transport pathways that diverted a large quantity of sediment from the paleo-Tennessee delta front to the adjacent slope toe and abyssal plain (Combellas-Bigott and Galloway 2002b). The McAVLU submarine fan system (named for the 3 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 their 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 a point source rather than a line source. Combined depositional loading and extension along the Gulf shelf margin caused continued compression along the Port Isabel and Perdido fold belts (Fig. 3.4B), triggered further shallow salt-canopy inflation beneath the paleo-continental slope of Louisiana, and initiated Miocene compression along the Mississippi fan fold belt (Fig. 3.2). The upper Miocene depositional episode records a long period (6 million years) of relative paleogeographic stability and high sediment supply (Galloway et al. 2000; Wu and Galloway 2002). Sediment input was dominated by the paleo-Mississippi and paleo-Tennessee systems. In the central Gulf, the shelf edge prograded 40 to 90 km (Fig. 3.9). 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 the northern Mexico to western Louisiana shelves. Accumulation in the east-central Gulf of up to 5 km of upper Miocene sediment continued basinward salt displacement and compression along the Mississippi fan fold belt. Along the wave-dominated northwestern Gulf, continental margin progradation onto muddy slope aprons built the shelf edge to or near its present position.

Early Pliocene Depositional Episodes The upper Miocene episode terminated with regional marine flooding associated with the last occurrence of benthic foraminifers Robulus E and / or Bigenerina A. The subsequent 2 million year depositional episode, which

is named for the contained Buliminella 1 faunal top, bridges the latest Miocene to early Pliocene (Fig. 3.8). Although clastic input continued to be focused through the paleo-Mississippi and Tennessee rivers, accumulation shifted onto the continental shelf and margin (Galloway et al. 2000) (Fig. 3.9). The McAVLU fan system was completely abandoned. Thickest deposits occur within a combined fluvialdominated delta system and subjacent delta-fed apron on the central Gulf margin. Upper slope minibasins captured the bulk of the sediment that spilled over the shelf edge (Winker and Booth 2000). Remobilization of the subjacent salt canopy is recorded in the South Timbalier Ship Shoal fault family of the subshelf roho domain (Fig. 3.2) (Schuster 1995). Beginning in the early Pliocene, rapid, high-amplitude, glacio-eustatic sea-level changes are manifested in the Gulf stratigraphic record by development of multiple sequences of 1 million years to several hundred thousand years duration with well-defined subaerial exposure and flooding surfaces (Lawless et al. 1997; Weimer et al. 1998; Prather 2000). Paleogeographic reconstructions and supply rate suggest these can be grouped into 2 composite genetic sequences (Fig. 3.9). Together with the Buliminella 1 episode, they constructed much of the eastern Texas and western Louisiana outer shelf (Fig. 3.9). Following the brief, post-Buliminella 1 transgression, the pattern of deposition changed in several ways (Galloway et al. 2000; Galloway 2005): 1. The paleo-Red River fluvial axis was rejuvenated by epeirogenic uplift and eastward tilting of the western High Plains and Rocky Mountains. 2. Sediment supply through the paleo-Tennessee River continued to decline. As a consequence, depocenters shifted to the northwestern Gulf margin, and the northeastern Gulf continental slope again became relatively sediment starved. There, Pliocene strata are thin, but locally sandy. 3. Shelf margin progradation occurred primarily along the west-central Gulf margin. 4. Along the relatively steep northeastern margin, turbidite-channel complexes extended to the slope toe, initiating a new submarine fan system informally called the WRLU fan for its location beneath Walker Ridge and Lund MMS areas. Deposition in this fan system continued for much of the remaining Pliocene. 5. Beneath the outer shelf, salt withdrawal caused

Geological Evolution ~ 47

active growth of the South Cameron fault family of the roho province (Fig. 3.2). On the slope, episodic depositional loading of the shallow salt canopy accentuated development of ponded-basin accommodation space, creating thick basin fills with numerous reservoir–seal pairs (Prather 2000).

Late Pliocene–Quaternary Depositional Episodes 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). As the valley was repeatedly cut and backfilled 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 (Saucier 1994). Development of a singularly large river draining into the central Gulf of Mexico created an extensive fluvial-dominated delta and subjacent slope apron system. Periodic, rapid transgressions forced shorelines temporarily landward 150 to 250 km, creating broad, short-lived, shallow shelves. Subsequent sea-level drawdown carved alluvial valleys across the shelf 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 frequent sea-level changes triggered a phase of mass wasting and submarine canyon erosion and filling unlike any previously seen in the basin (Prather 2000). Canyon excavation was most active on the eastern flank of the delta systems. 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 minor 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 downslope from prograding shelf-margin deltas. Salt mobilization and loading beneath the outer shelf is recorded by growth of the roho domains (Fig. 3.2).

Summary Depositional history of the northern Gulf of Mexico reflects and amplifies the story of tectonism within the North American continent. The tectonically modulated composite episodes exert first-order control on continental margin offlap and bypass of sand to the basin center. Five major phases of continental uplift and erosion are recorded in the shifting patterns and rates of supply (Galloway 2005). (1) 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. (2) Paleocene through middle Eocene pulses of Laramide uplift along the central and southern Rocky Mountains and Sierra Madre Oriental supported the early Cenozoic depositional episodes. Although the prograding shelf and slope lay primarily beneath the modern coastal plain (Fig. 3.9), drilling and major petroleum discoveries on the Texas–Louisiana continental slope demonstrate that abundant sand-rich sediment bypassed the margin far onto the Gulf abyssal plain. (3)  From the 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 depositional episodes from the Oligocene through early Miocene. Resultant Frio offlap extended the paleo-shelf to the present Texas coastline. (4) Miocene exhumation of the Cumberland Plateau and Appalachians invigorated supply to the east-central Gulf basin. At the same time, the Rocky Mountain uplands experienced continued regional erosion. Miocene offlap constructed much of the Texas shelf and initiated outbuilding of the Louisiana shelf (Fig. 3.9). (5) Pliocene uplift of the High Plains further rejuvenated Rocky Mountain sources and created a broad eastward slope that converged with the west-sloping alluvial apron of the eastern interior. East- and west-flowing streams were variously combined and directed southward, forming the distinct Red and Mississippi fluvial axes that dominated Plio-Pleistocene sedimentation in the west-central Gulf shelf. High rates of Pleistocene sediment accumulation reflect rapid Quaternary climate cycling and glacial erosion and runoff directly into the principal sediment transport systems. The central Gulf shelf was prograded to its current position (Fig. 3.9). At the same time, increasingly high-amplitude and high-frequency glacio-eustatic sealevel changes have further enhanced the natural instability of the continental margin.

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Conclusions and Research Directions The grand theme that emerges from the recent history of the northern Gulf of Mexico is one of “mobility”: 1. The mobility of salt and thick, under-consolidated sediment beneath the spatially and temporally variable sediment load has resulted in wholesale transfer of mass from basin margin toward the basin center, underpinning and even upbuilding the continental shelf and slope (Rowan 2002). Evacuation of salt has created space for thick, local-to-subregional sediment accumulations, exerting an arguably dominant control on location of accommodation volume. 2. The mobility of sediment in response to gravity transport processes has distributed sand from the shelf margins downslope and far across the basin floor. The known volume, extent, and stratigraphic distribution of variously described “deepwater,” “lowstand,” or “turbidite” reservoirs continues to grow and repeatedly make obsolete the presumed limits to reservoir development in deep marine basins. In the Paleocene, bypass was particularly effective as Wilcox deltas built onto a relatively steep, carbonate-constructed upper slope that was minimally complicated by salt-canopycreated topography. Likely, a southward slope of the Paleocene abyssal plain, reflecting the broad crustal transition inherited from asymmetric rifting, would have aided gravity-flow transport far into the basin (Pindell et al. 2003). Neogene abyssal submarine fan systems record active bypass that was concentrated along structural alignments or large erosional canyons located along the relatively steep and graded northeastern Gulf of Mexico slope. 3. The mobility of contained fluids is reflected in the extent and multiplicity of hydrocarbon systems identified in the northern Gulf. Source rocks beneath the modern continental shelf and slope are found primarily in basinal facies of the Upper Jurassic Smackover and Cotton Valley episodes, and Paleocene–Eocene episodes (Hood et al. 2002). These strata are buried far below the dominant reservoir-bearing Miocene, Oligocene, and Plio-Pleistocene sequences. Thus, the center of mass of fluid hydrocarbons has migrated kilome-

ters to tens of kilometers through space and tens of millions of years through geologic time. The Gulf basin is, in many ways, one of the best understood basins in the world. At the same time, it displays a geologic complexity and diversity that continues to pose new geologic questions; open new venues for renovating, updating, and, sometimes, challenging existing paradigms; and it provides a world-class exploration province for the 21st century. Current research and exploration themes include: 1. Imaging and interpreting subcanopy and subweld structural and sedimentary features beneath the continental shelf and upper slope. 2. Filling in the large data gap between updip Paleocene–Miocene continental platform strata as seen by exploration beneath the coastal plain and their abyssal equivalents that re-emerge at drillable depths in the lower continental slope. 3. Developing the necessary models and computational tools for the three-dimensional reconstruction of evolving Gulf margin subsidence, morphology, and structural growth history. 4. Quantitatively assessing and predicting the effects of geochemical, thermal, and hydrological processes that control sandstone consolidation and reservoir quality at depths exceeding 6 km.

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Plio-Pleistocene sequence boundaries, Gulf of Mexico: oxygen isotopes, ice volume, and sea level. Pp. 495–98 in J. M. Armentrout and N. C. Rosen (eds.), Sequence Stratigraphic Models for Exploration and Production: Evolving Methodology, Emerging Models and Application Histories: 22nd Annual GCSSEPM Foundation Bob F. Perkins Research Conference. Houston, Tex.: GCSSEPM Foundation. Marton, G., and R. T. Buffler. 1993. Application of simple-shear model to the evolution of passive continental margins of the Gulf of Mexico basin. Geology 21:495–98. ———. 1999. Jurassic–Early Cretaceous tectonopaleogeographic evolution of the southeastern Gulf of Mexico basin. Pp. 63–91 in P. Mann (ed.), Caribbean Basins. Sedimentary Basins of the World, 4. Amsterdam: Elsevier Science. McFarlan, E. Jr., and L. S. Menes. 1991. Lower Cretaceous. Pp. 181–204 in A. Salvador (ed.), The Geology of North America, Volume J, The Gulf of Mexico Basin. Boulder, Colo.: Geological Society of America. Mitchum, R. M., J. B. Sangree, P. R. Vail, and W. W. Wornardt. 1993. Recognizing sequences and systems tracts from well logs, seismic data, and biostratigraphy: examples from the late Cenozoic of the Gulf of Mexico. Pp. 163–98 in P. Weimer and H. Posamentier (eds.), Siliciclastic Sequence Stratigraphy. Memoir 58. Tulsa, Okla.: American Association of Petroleum Geologists. Nehring, R. N. 1991. Oil and gas resources. Pp. 445-94 in A. Salvador (ed.), The Geology of North America, Volume J, The Gulf of Mexico Basin. Boulder, Colo.: Geological Society of America. Peel, F. J., C. J. Travis, and J. R. Hossack. 1995. Genetic structural provinces and salt tectonics of the Cenozoic offshore U.S. Gulf of Mexico. Pp. 153–17 in M. P. A. Jackson, D. G. Roberts, and S. Snelson (eds.), Salt Tectonics: A Global Perspective. Memoir 65. Tulsa, Okla.: American Association of Petroleum Geologists. Pindell, J. L., and L. Kennan. 2001. Kinematic evolution of the Gulf of Mexico and Caribbean. Pp. 193–220 in R.H. Fillon, N.C. Rosen, and P. Weimer (eds.), Petroleum Systems of Deep-Water Basins: Global and Gulf of Mexico Experience: 21st Annual GCSSEPM Foundation Bob F. Perkins Research Conference. Houston, Tex.: GCSSEPM Foundation. Pindell, J. L., L. Kennan, and T. Watts. 2003. Asymmetric rifting and the northern Gulf of Mexico supra-salt platform: implications for the initial depositional setting of Texas– Louisiana Tertiary clastic systems. Pp. 135–38 in in H. R. Roberts, N. C. Rosen, R. F. Fillon, and J. B. Anderson (eds.), Shelf Margin Deltas and Linked Down Slope Petroleum Systems: Gobal Significance and Future Exploration Potential

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␥4

Northern Gulf of Mexico Sea-Level History for the Past 20,000 Years James H. Balsillie and Joseph F. Donoghue

High-resolution, composite sea-level curves have been developed for the northern Gulf of Mexico for the period since the Last Glacial Maximum. The goal of this work was twofold: (1) to define the regional sea-level history of the northern Gulf of Mexico using all of the available geochronological data on sea-level history, and (2) to examine the hypothesis that, for stable coastal regions of the Gulf of Mexico coastline, sea-level history approximates global (i.e., eustatic) sea level. The resulting sea-level curves are based on all available carbon-dated indicators of paleo-sea level and represent, on average, one measurement every 65 years for the past 20,000 years. The data sets consist primarily of geological sea-level indicators, along with some dates from archaeological artifacts. Published sea-level histories for the Gulf of Mexico exhibit significant variability. While error is associated with the 14C dating methodology, most error is undoubtedly associated with the choice of sample and sampling location representing paleo-sea level. The present analysis is the first exhaustive treatment of such error for a comprehensive Gulf of Mexico sea-level data set. The data set described here represents a compilation and analysis of all of the available and usable sea-level data sets for the northern Gulf of Mexico. For each of the data subsets, the associated error was analyzed and the data were edited using 3 procedures (one geologi-

cal, and the other 2 statistical) to construct a sea-level curve which is as accurate as the data permit. First, we selected data for all regions of the northern Gulf of Mexico exhibiting crustal stability. This process yielded 353 radiocarbon-dated sea-level indicator data points. Second, we addressed the problem of identifying sea-level data outliers that can justifiably be excluded in analytical procedures. This is not, in fact, a problem peculiar to Gulf of Mexico data, but rather is the typical case for most sea-level data sets. Using a recent global eustatic data set from the Red Sea as a basis for comparison, a statistically based method has been developed to provide a useful tool for the editing of sea-level data in general. We found that only a few outlier data points can significantly affect analytical outcomes (only 12 outliers, representing 3.4% of the data set, were identified among the Gulf of Mexico sea-level indicators). Third, once such outliers were eliminated, a sufficient amount of data still remained (341 dated sea-level indicators) with some associated variability. We treated the remaining data using a 7-point floating average method. By smoothing some of the noise, the moving average method removed much of the variability while allowing longer-term trends to remain. On average, 7 dated points encompass a period of approximately 400 years, with each floating-point average representing a 65year period.

This chapter is in memory of Jim Balsillie who passed away after a long illness. He was a pioneer in coastal sedimentary research and added new dimensions to our understanding of coastal change. His many admirers in both the geologic and engineering communities will miss his sharp insight and good humor. He was a good geologist and great friend. 53

54 ~ Balsillie and Donoghue

Additionally, we investigated the controversial subject of sea-level history younger than about 6000 years (mid- to late Holocene) and established 2 “younger data sets” based on sampling location bias. One younger data set can be defined by sea-level indicators collected exclusively seaward of the present shoreline (younger data set A) and the other by sea-level indicators collected landward of or near the present shoreline (younger data set B). By definition, sea-level data sampled seaward of current sea level exclude potential highstand shoreline samples. In addition, a discussion of littoral processes associated with physiographic features (beach ridges, cheniers, and storm ridges) has been included for the purpose of justifying inclusion of the data in younger data set B. The resulting comprehensive compilation of northern Gulf of Mexico sea-level history has significance beyond the Gulf region. Gulf of Mexico data compare favorably with a recent late Quaternary sea-level data set from the Red Sea, a high-resolution index of eustatic sea level. Given its geologic stability throughout the late Quaternary, its distance from ice sheets, and its relatively lowenergy environments, the northern Gulf of Mexico might be expected to have experienced near-eustatic sea-level conditions. The northern Gulf of Mexico sea-level data therefore embody a detailed record of global sea level. In particular, the strong evidence of mid- to late-Holocene sea-level highstands in the northern Gulf of Mexico may be among the best global verifications of such events.

Radiocarbon Versus Absolute Dates All Gulf of Mexico sea-level data in the data sets we used are based on radiocarbon dating of paleo-shoreline indicators. A variety of analytical problems can affect radiocarbon age determinations. Radiocarbon ages are given in years before present (yr BP, where present is taken to be 1950), with a plus-and-minus error representing the standard deviation. One of the assumptions made in radiocarbon dating is that the 14C content of the dated sample has not changed after the death of the organism other than change due to radioactive decay. This assumption is not always fulfilled, as documented by Mook and van de Plassche (1986). An additional source of radiocarbon dating error concerns the 14C half-life. By long-term convention, the 14C half-life used in age determinations is 5568 years. This value is in error by 3% and should be 5730 years. Whether or not published data sets have been corrected for this discrepancy may not be apparent.

Assuming that, in published results, such problems as those above have been corrected to the maximum extent possible, 14C dates still do not represent true calendar years. Radiocarbon years would be equivalent to calendar years only if the 14C concentration in the atmosphere were constant over time. This has been shown not to be the case. Atmospheric 14C concentration has fluctuated over time due to variation in cosmic radiation intensity, fossil fuel burning, and nuclear testing (Faure 1986; Suess 1986). To understand sea-level change in terms of absolute or sidereal time, radiocarbon dates for the current data set can be converted using a calibration scheme. Radiocarbon calibration methods are based on comparing radiocarbon dates with actual ages for samples whose absolute age has been determined independently, such as via tree rings or lake varves, or in some cases by other radiometric dating techniques. One of the standard calibration schemes, incorporating dendrochronologically dated wood samples, is the CALIB program produced by the Quaternary Isotope Laboratory of the University of Washington (Stuiver and Kra 1986; Stuiver and Reimer 1993; Stuiver et al. 1998a, Stuiver et al. 1998b; McCormac et al. 2002). Several calibration data sets are available. For terrestrial materials, the IntCa198 decadal data set (1998 atmospheric delta 14 C; Stuiver et al. 1998a) can be applied to data from the Gulf of Mexico region. For marine material, the Marine98 data set (1998 marine delta 14C; Stuiver et al. 1998b) can be used where regional offsets can be applied (e.g., Stuiver and Braziunas 1993; Stuiver et al. 1998b). This application, along with any regional offsets, appears to provide the best calibration available. Using CALIB (Rev 4.4.2), the current Gulf of Mexico data were converted to absolute years and reported as calibrated years before present (cal yr BP). Only the calibrated age data will be shown in this chapter. The full data set, in both radiocarbon and calibrated years, is available in Balsillie and Donoghue (2004).

A Global Sea-Level Record from the Red Sea We begin the analysis of Gulf of Mexico sea-level history by examining a recent effort to develop a “eustatic” sealevel record for the late Quaternary. Siddall et al. (2003) and Siddall et al. (2004) presented an original method for determining global sea-level changes for the last glacial cycle, using δ18O analyses of foraminiferal sediments from a Red Sea sediment core. This method has met with con-

Sea-Level History ~ 55

siderable interest as a new approach to defining eustatic sea-level change (e.g., Rohling et al. 2003; Sirocko 2003). The geomorphology and hydrology of the Red Sea basin, combined with effects due to low latitudes, make the Red Sea δ18O data a sensitive indicator of global sealevel change. Low latitudes equate to high evaporation rates, leading to higher salinities for ocean water bodies and, hence, enriched 18O levels. For the Red Sea the only significant link with oceanic waters is the southern entrance, the strait of Bab el Mandab, which is only 18 km wide. Furthermore, a sill at the entrance restricts water flow. At present sea level, the top of the sill lies at a depth of about 137 m below mean sea level. During the Last Glacial Maximum, it lay at a depth of only about 15 m below mean sea level. At low sea-level stands evaporation and increased salinity resulted in stronger δ18O signatures. As a consequence, Red Sea sediments contain a greatly amplified δ18O record for progressively lower sealevel stands. Siddall et al. (2003) constructed a numerical model for correlating δ18O data with sea level and tied the results to 5 14C age markers. The significance of the Red Sea work lies in how it relates to the δ18O record from polar ice cores. The Byrd and Vostok ice cores from Antarctica and the GISP2 ice core from Greenland have been correlated with the Red Sea data. The temperature records obtained from oxygen-isotope data from high-latitude ice cores are directly correlated with a high-resolution marine record of sea-level change (Sirocko 2003). In addition to δ18O data from Red Sea foraminifera, Siddall et al. (2003) also included 14C coral data from Barbados (Fairbanks 1989, 1990; Bard et  al. 1996), Tahiti (Bard et al. 1996), and New Guinea (Edwards et al. 1993) to augment their global eustatic sea-level history (Table 4.1). We present these global sea-level curve data (plotted in Fig. 4.1) because they represent a reference for identifying outliers in regional data such as our Gulf of Mexico data sets. A representative transcendental equation has been fitted to the global sea-level data (i.e., the 7-point floating average curve of Figure 4.1). The equation and plotted results are shown in Figure 4.2.

Northern Gulf of Mexico Composite Sea-Level Curve Twenty-three sea-level data sources or subsets for the northern Gulf of Mexico region (Table 4.1) were examined, resulting in 353 dated sea-level indicators. The data

cover approximately the past 20,000 years of geologic time. The data are plotted in Figure 4.3, but not all data subsets are plotted. Data younger than about 7000  cal yr BP, if plotted at small scale, would render the figure illegible due to the high concentration of data points. Therefore, the data are divided into 2 age ranges: (1) ages between about 20,000 and 7000 cal yr BP, and (2) ages younger than 7000 cal yr BP.

Identifying Data Outliers As mentioned above, we sought all sea-level data sets from stable regions of the Gulf of Mexico. Certain regions of the Gulf do not meet that criterion. The Mississippi Delta, with its well-documented subsidence history, includes many sea-level data sets which must be judiciously examined for use in a regional sea-level history. For example, the majority of McFarlan’s (1961) Mississippi Delta data were excluded, with the exception of the younger beach and chenier data (6 and 10 m deep line the seaward edge of the main bedrock depression beneath Hawk Channel. The superposed data definitively placed the mid-channel patch reefs along the landward margins of both troughs. The troughs and higher topography of the outer shelf have shielded the patch reefs from surf and outer-shelf sediment transported landward by storms. The troughs are presumed to consist of non-coralline grainstone and packstone such as that cored beneath Hawk Channel off Key Largo (Shinn et  al. 1994; Multer et  al. 2002). The shallower trough is off the Upper Keys. The deeper trough is off the Middle and Lower keys, consistent with deeper bedrock topography to the southwest. Contoured as topographic lows with thicker sediments than surrounding higher-elevation areas, neither was defined as an enclosed entity on the updated

Table 18.2. Data and assumptions used to infer youngest possible age of upper-slope terrace off the Florida Keys (modified from Lidz 2004). Without reference to any particular coral species or growth rate, thickness (~8 m) of the post- Q3–Unit section in core 56 (Perkins 1977) indicates that a comparable 7.3 m of Stage- 6 / 5 (substage-5e) coral framework could have accreted at the base of the outlier reefs. That said, the upper-slope terrace might be older than 190 ka. The only way to verify its age is to core and date the material at and below its surface, an unlikely procedure given the water depths and strong currents of the offshore setting.

Material dated

Dates (ka)

Geologic age

A. palmata, C. natans

8.9–6.9

Holocene

1 A. palmata, 1 C. natans, 8 M. annularis

84.5–80.9

5a

3 M. annularis, 1 A. palmata

94.4–90.6

5b

M. annularis

106.5

5c

M. annularis

128.1–112.4

5e

M. annularis, top of Q5 Unit

(highstand) Marine Q4 Unit top

~125

140 mid-Pleistocene

‡

‡

‡

‡

‡

5e

Coral depths † (m)

Thickness in outlier reef (m)

Maximum highstand † (m)

Authors

–12.3 to –8.9

3.4

0

Toscano and Lundberg 1998

–24.0 to –12.3

11.7

–9

Ludwig et al. 1996; Toscano and Lundberg 1999

–19.8 to –15.5

4.3

–21.7

?

–36.6 to –15.9 ~ +5.5

not identified in Florida –5.0

–14 to –10

Toscano and Lundberg 1999

–15

Toscano and Lundberg 1999

~ +10.6

Multer et al. 2002

~ +10.6

Hoffmeister and Multer 1968; Perkins 1977; Halley and Evans, 1983 Chappell 1974 Muhs 2002; Muhs et al. 2004 Chappell 1974

(highstand)

185

(highstand)

220

?

Marine Q3 Unit top

~366.8

–8.0

Multer et al. 2002; Muhs et al. 2004

Marine Q2 Unit top

>366.8

–37.5

Multer et al. 2002

not identified in Florida

Chappell 1974

Notes: Coral dates and depths of Toscano (1996) and Toscano and Lundberg (1998, 1999) are from the largest Sand Key outlier reef. Coral dates of other authors are from elsewhere in the Florida Keys. †

Depths relative to present sea level. Depths of Q2–Q4 unconformities from Big Pine Key core 56 of Perkins (1977). Elevation for top of Q5 Unit is the highest elevation in the keys (at Windley Key; Stanley 1966; Lidz and Shinn 1991). ‡

Substages of marine oxygen-isotope Stage 5.

Data:


t
6QQFSTMPQF
UFSSBDF
PO
XIJDI
UIF
PVUMJFS
SFFGT
HSFX
JT
o
N
CFMPX
QSFTFOU
TFB
MFWFM


t
-BSHFTU
4BOE
,FZ
PVUMJFS
SFFG
JT
_
UP

N
NFBO

N
JO
TFJTNJD
SFMJFG


t
ćF
PMEFTU
DPSBMT
EBUFE

LB
BSF
BMTP
UIF
EFFQFTU

N
CFMPX
TFB
MFWFM
DPSFE
GSPN
UIF
PVUMJFS
SFFG
XIJDI
MFBWFT
_
N of corals of unknown age at the base of the reef.


t
ćF
OFBSFTU
TVCTVSGBDF
EBUB
UP
UIF
4BOE
,FZ
3FFG
BSFB
DPNF
GSPN
B
MPOH
DPSF
DPSF

BU
_
N
SFDPWFSFE
GSPN
#JH
1JOF Key (Perkins 1977). Unconformities on top of the Q1 and Q2 units in core 56 are found at respective depths of ~48.0 and 37.5 m





CFMPX
TFB
MFWFM
ćFTF
EFQUIT
BSF
CFMPX
FMFWBUJPO
PG
UIF
UFSSBDF



t
4JY
TUBOET
PG
1MFJTUPDFOF
TFB
MFWFM
EVSJOH
UIF
QBTU

UIPVTBOE
ZFBST
XFSF
IJHI
FOPVHI
UP
IBWF
ĘPPEFE
UIF
UFSSBDF
BOE
TIFMG
(Imbrie et al. 1984). Assumptions:


t
ćF
UFSSBDF
VOEFS
UIF
PVUMJFS
SFFGT
QSFEBUFT
UIF
SFFGT



t
ćF
UFSSBDF
JT
FSPTJPOBM
BOE
XBT
GPSNFE
EVSJOH
B
GBMMJOH
TFB
MFWFM



t
"HF
PG
UIF
CPUUPN

N
PG
UIF
PVUMJFS
SFFG
JT
PMEFS
UIBO

LB



t
ćF
DMPTFTU
QPTTJCMF
UJNF
UP

LB
GPS
PMEFS
DPSBM
HSPXUI
PO
UIF
UFSSBDF
JT
EVSJOH
UIF
NBSJOFJTPUPQF
TVCTUBHFF
IJHITUBOE
PS







EVSJOH
B
TMJHIUMZ
PMEFS
UJNF
BU
_
LB
BT
B
SJTJOH
TFB
ĘPPEFE
UIF
UFSSBDF
EVSJOH
UIF
JTPUPQF
4UBHF
USBOTJUJPO



t
1SJPS
UP

LB
UIF
BQFY
PG
UIF
MBTU
TUBOE
PG
TFB
MFWFM
UP
ĘPPE
UIF
UFSSBDF
BOE
TIFMG
PDDVSSFE
BU
_
LB
'JH

*NCSJF
FU
BM






 
ćF
NBSJOF
SFHSFTTJPO
GSPN
UIBU
IJHITUBOE
SFBDIFE
UFSSBDF
EFQUIT
BU
_
LB
UIF
NPTU
SFDFOU
UJNF
QPTTJCMF
UIBU
UIF
terrace could have been formed. If this assumption is true, then the terrace is a regressional erosional feature.

314 ~ Lidz

0

Pleistocene backreef trough

30

multiple of Holocene surface

45

Depth (m) below sea level

15

Line 3 multiple of Pleistocene surface

A

60

NNW

SSE

sea level

0

Pelican Shoal Straits of Florida horizontal Holocene layers

poorly developed reef or PALEOSHORELINE hardground ridge?

Pleistocene shelf-margin reef

15

PALEOSHORELINE

backreef trough at ~23-m depth

Holocene sediment

concave bedrock surface

30

Pleistocene terrace

inclined Holocene layers

45

Depth (m) below sea level

~7 m

Holocene infill

Pleistocene surface 1 km vertical exaggeration 15x

B

60

24°29.7'N 81°37.4'W 1650

24°29.0' 81°37.2' 1700

24°29.3' 81°37.3' 1655

24°28.7' 81°37.2' 1705

GPS coordinates and hour

Figure 18.7. (A) Seismic profile obtained in 1989 and (B) interpretation show upper-slope terrace off Pelican Shoal (Lower Keys, Figs. 18.1C, 18.3B) (from Lidz et al. 2003). Profile crosses shelf margin nearly normal to a ~20-km–long, single outlier-reef tract east of the multiple outliers off Sand Key Reef. Note facies change in the sediment wedge on the terrace. Horizontal layers are interpreted to result from a stillstand that eroded tops of the inclined strata and redeposited the sands in an intertidal zone. Latitude and longitude are in decimal minutes based on GPS coordinates. Hours adjacent to ticks are navigational correlation points along the seismic line.

bedrock-surface or Holocene isopach maps (Lidz et al. 2003). Those maps were constructed without benefit of the later habitat and NGDC sets of data. Neither trough is evident in the aerial-photo or habitat data alone (Figs. 18.3A, B). Both troughs harbor the seagrass and limemud habitat characteristic of the seabed beneath Hawk Channel (Lidz et al. 2006). These are the first such nuclei for corals—inner-shelf trough edge, likely grainstone and packstone substrate—known in the Florida reef record.

Two Coral Architectures and Asymmetrical Evolution The south Florida shelf built upward and seaward through repetitive accretion of 2 principal, discontinuous coral architectures: narrow ridge-and-swale and broad reefand-trough structures that produced geometries of very different scales. Isolated outer-shelf coral-rock ridges are well known to local residents, but ridge number, spatial continuum, and geologic significance were not recognized

Coral Reefs of the Florida Keys ~ 315

15

schooling fish

30

multiple of Pleistocene backreef trough 45

Depth (m) below sea level

0

multiple of Holocene surface 60

Line 12

A

Maryland Shoal

S

sea level

Holocene ~7 m

Pleistocene shelf-margin reef buried outlier reef

backreef trough at ~ 18-m depth

0

Straits of Florida 15

outlier reef Holocene sediment

30

Holocene sediment Pleistocene terrace 45

Pleistocene surface 60

B

1 km vertical exaggeration 22x 24°29.8'N 81°35.8'W 1815

24°29.4' 81°35.7' 1810

24°29.1' 81°35.8' 1806

GPS coordinates and hour

before assembling the mosaics needed for mapping habitats. A ridge-and-swale couplet is several tens of meters across. The ridges are below seismic resolution (30 m) from the center and base of the mas-

316 ~ Lidz

sive reefs. The logistics of drilling, and penetrating the base of an outlier reef, in 12 m of water in the strong currents of an open-marine setting are too great. These 2 reef structures form the basis for the geomorphogenic models of shelf-edge stratigraphy within 40 m of water (Lidz 2004). The sites modeled (Carysfort Reef, The Elbow, Pelican Shoal, and Sand Key Reef) were selected because of the presence of coral reef complexes, geomorphic dissimilarities, and quantity and quality of data at the sites. A “generic” profile for the model “floor” was based on the simple seismic record of the upper-slope terrace and dune-ridge facies at Pelican Shoal (Figs. 18.1C, 18.7A). Reflections representing the Pleistocene surfaces were traced from seismic profiles and affixed to the drawings. “Layers” simulating Pleistocene coralline stratigraphic sections were added between the terrace and Pleistocene surfaces. Layer thickness was arbitrary due to model scale. The layers were correlated with coral dates at the sites and with interglacial sea-level maxima (Figs. 18.4A, B, Table 18.2). Although the tops of the coralline Q3 and Q4 units at the west end of Key West are ~21 and 10 m below present sea level, respectively (Multer et al. 2002), layers representing Q3 and Q4 times (~340–300 ka and ~230–220 ka; Muhs et al. 2004) are not present in the models because Q3 and Q4 deposition pre-dated the inferred minimum age of the terrace, ~190 ka. The traced reflection representing the Holocene surface (complete stratigraphic section) was added on top. Carysfort Reef and The Elbow are Holocene buildups located ~9 km apart on the Pleistocene shelf-edge reef off the Upper Keys (Fig. 18.1C). Two main geomorphic differences distinguish the sites: (1) A prominent outlier reef exists seaward of Carysfort Reef (Fig. 18.9) but not off The Elbow (Fig. 18.10). On the basis of differences in scale, the outlier is considered to represent a tract different from those of the outlier reefs off Sand Key Reef (Lidz et al. 1997b; Toscano and Lundberg 1999; Lidz et al. 2003). (2) The trough behind the Carysfort outlier is only partially filled with Holocene sediments, whereas the trough behind the shelf-edge reef under The Elbow is completely filled. Both areas have an elevated bedrock ridge at the seaward edge of the upper-slope terrace. Both areas display buried low-relief terrace accretions that are acoustically impenetrable and linear, as is evident from other seismic profiles across the Upper Keys margin (Lidz et al. 1997b). As interpreted from seismic facies, the accretions represent 4 parallel coral-capped ridges that may have been incipient reefs equivalent to the outlier tracts off the Lower Keys (Lidz et al. 1997b; Lidz et al. 2003). According

to the models, the Pleistocene sequences on the Carysfort outlier and at The Elbow have accreted upward and landward, indicating that Pleistocene corals may have backstepped through time. At least one instance where Pleistocene corals backstepped ~0.25 km between 85.3 and 80.2 ka has been confirmed with coral dates from a landward-oriented core transect on the Carysfort outlier (Toscano and Lundberg 1998). As evidenced by several other dates on the outlier, Holocene coral growth advanced seaward over more than 0.5 km before backstepping toward the end of accrual (Toscano and Lundberg 1998). From the Carysfort model, the Pleistocene shelf-edge reef was also shown to have seaward-accreting components. Pelican Shoal and Sand Key Reef are Holocene buildups ~26 km apart on the shelf-edge reef off the Lower Keys (Fig. 18.1C). These sites also have numerous, very different geomorphic and stratigraphic characteristics. The terrace surface off Pelican Shoal has remained essentially unchanged since it formed (Figs. 18.7A, 18.11), whereas 4 massive high-relief tracts of outlier reefs grew on the terrace off Sand Key Reef (Fig. 18.12). A sharp seismic stratigraphic contact at a water depth of ~25 m divides inclined and flat bedding in a Holocene sediment wedge off Pelican Shoal (Fig. 18.7A). No such contact is detected in sediments between the Sand Key outlier reefs. As evident from seismic facies, a rounded topographic high (beach-dune ridge?) caps the seaward edge of the terrace off Pelican Shoal. The ridge served as the nucleus for accretion of the largest outlier reef off Sand Key Reef. The seismic reflection of the ridge is analogous to those of the deep dune ridges off the Marquesas Keys (Locker et al. 1996; and Hine and Locker in chapter 7 in this volume). The trough behind Pelican Shoal is V-shaped but is U-shaped behind Sand Key Reef (Figs. 18.11, 18.12). Despite alongshore sand transport in the Lower Keys, both trough areas on the shelf are filled with ~16 m of storm-generated Holocene coral rubble and sand. Clearly evident in aerial photographs as elongate debris fields, reef-rubble zones extend landward behind most major shelf-edge reefs, obscuring margin-parallel seabed features (Figs. 18.3A, B, 18.6B). Field observations revealed landward-dipping bedding in the rubble (Ball et al. 1967; Enos 1977). Consistent with increased current intensity and off-bank sediment transport due to the open-marine setting, the deeper, wider troughs behind the Sand Key outlier reefs are not yet filled. Irregularities in evolution of the shelf-edge reef were minor and similar at the 4 sites modeled. Pleistocene

Florida Keys Key Largo Ls

Hawk Channel

Holocene lime mud

Carysfort outlier reef

outer shelf

?

?

Holocene ?

Straits of Florida

sea level

77.8 ka† 7.6-5.0 ka‡ 85.3-80.2 ka§ 94.4-92.2 ka§ Holocene sediments

shelfmargin reef

†Multer et al., 2002 ‡Toscano and Lundberg, 1998 §Toscano, 1996

outlier reef Pleistocene surfaces ~ 1 km

Key Largo Ls

vertical exaggeration 10x

sea level

highstands at ~105, 90, and 80 ka† (Q5 Unit‡, isotope substages-5c, 5b, and 5a§ time)

?

Hawk Channel

†Chappell, 1974 ‡Perkins, 1977 §Chappell and Shackleton, 1986

shelfmargin reef

outlier reef

coral growth

sea level ~125-ka highstand (Q5 Unit, isotope substage-5e time)

TIME

Key Largo Ls

Hawk Channel coral growth

coral growth

sea level ~127-ka transgression (Stage-6/5 transition)

Hawk Channel lagoon Imbrie et al., 1984

initial coral growth at margin and on terrace Hawk Channel depression

beach-dune ridge deposits shelf exposed

Pleistocene surface upper-slope terrace (~190 ka) post-195-ka lowstand

sea level ~195-ka highstand

Hawk Channel lagoon

NW

initial coral growth on beach-dune ridges

inferred paleotopography

SE

Figure 18.9. Geomorphogenic model of Carysfort Reef shows a Pleistocene backstepped reef complex in the outlier reef and Holocene backfilled progradation of the outer shelf (top panel). The shallow Holocene surface at and behind the shelf edge in this model is drawn from field knowledge. Though beach-dune ridges are suspected to underlie all coral reefs shelfwide, they are discontinuous; their presence behind Carysfort Reef has not been confirmed. Note shapes of the outlier reef and its landward trough are very different from those of the Sand Key Reef outliers (Fig. 18.12). Also note buried reef-like features on the upper-slope terrace. In all 4 models (Figs. 18.9–18.12), timing of pre-125-ka events is revised from Lidz (2004) based on the δ18O record (Fig. 18.4A). Horizontal scale and vertical exaggeration, shown for present stratigraphies (top panels), apply to features at and seaward of the shelf margin based on seismic profiles in Lidz et al. (2003). Shelf features are based on cores at Marker G (Figs. 18.2B, 18.3A) and are not drawn to scale. Long curved arrows show off-bank sediment transport. Short curved arrows show landward sediment transport and infilling of back-reef troughs. Up arrows denote 6 of the 7 highstands that produced coral reef complexes (Table 18.1). Dotted lines signify Pleistocene stratigraphic “layers.”

318 ~ Lidz

Florida Keys Key Largo Ls

Hawk Channel

outer shelf Holocene coral ridges and sands

The Elbow

Straits of Florida

sea level

Holocene

lime mud coral ridges and cemented grainstone (~80 ka)

Holocene sediments

shelfmargin reef Pleistocene surfaces ~ 600 m

Key Largo Ls

vertical exaggeration 7x

sea level

highstands at ~105, 90, and 80 ka† (Q5 Unit‡, isotope substages-5c, 5b, and 5a§ time)

Hawk Channel coral-ridge growth with sand-filled swales †Chappell, 1974 ‡Perkins, 1977 §Chappell and Shackleton, 1986

shelfmargin reef

sea level ~125-ka highstand (Q5 Unit, isotope substage-5e time)

TIME

Key Largo Ls

Hawk Channel coral-ridge growth with sand-filled swales

coral growth

coral growth

Pleistocene surface upper-slope terrace (~190 ka) Post-195-ka lowstand and 127-ka transgression phases are similar to those shown in Figure 9. There may be several layers of outer-shelf coral ridges and infilled swales of discrete ages beneath the Holocene layer. sea level Hawk Channel lagoon initial coral growth on beach-dune ridges

~195-ka highstand

inferred paleotopography

NW

highstand progradation potential had previously existed to the extent that backstepping reefs were slowly filling landward troughs, but the time needed (i.e., duration of transgressive intervals) was insufficient and the coral-infill process at those sites was uneven and incomplete. Coral dates from 3 Sand Key outlier-reef cores are evidence that backstepping took place at the core sites after ~82.7 ka—at 81.6 and at 80.9 ka (Toscano and Lundberg 1998). At each of the modeled sites, the Pleistocene component represents a backstepped reef-complex margin (Lidz 2004) even though coral growth had not completely closed the troughs. In contrast to the minor, but similar, irregularities in evolution of the shelf-edge reef is evidence for vastly dissimilar reef accrual on the upper-slope terrace. Seismic profiles across the terrace off Pelican Shoal (Fig. 18.7A) and off Sand Key Reef (Fig. 18.6B) constitute end-member

SE

Figure 18.10. Geomorphogenic model of The Elbow shows a Pleistocene backstepped reef complex at the margin and Holocene backfilled progradation of the outer shelf (top panel). Note the presence of coral-ridge bands and sediment-filled swales. Storm-generated bedding in Holocene fill behind The Elbow dips landward. Angle of repose at debris surface is ~45° (E. A. Shinn, University of South Florida, personal communication, 2003). Buried terrace outlier reefs are ~5–12 m high. Their crests are ~35–38 m below sea level. Compare with dimensions of highrelief outlier reefs at Sand Key Reef (Fig. 18.12).

views of reef stratigraphy and are evidence that asymmetrical evolution also occurred on a large and wide-ranging scale. Countless differences exist in dimension, discontinuity, number, and relief among the tracts of outlier reefs. The only common threads are terrace location and reef orientation parallel to the margin. The 4 freestanding outlier reefs off the Lower Keys (Lidz et al. 2003) (Figs. 18.6A, B) display keep-up profiles with 30 m of relief (Neumann and Macintyre 1985). The 4 coral-capped seismic facies believed to be their counterparts off the Upper Keys (Lidz et al. 1997b) (Fig. 18.10) display low-relief give-up (buried) profiles. In between, areas of 0, 1, 2, or 3 tracts are found in seismic profiles. Why the Upper Keys outliers did not keep up with, or were unable to catch up to, rising sea level is not known, but their inability to grow, in direct contrast to that of the Lower Keys outliers, is proof of the inconsistent responses of corals and reefs to rising sea level.

Coral Reefs of the Florida Keys ~ 319

Florida Keys

Hawk Channel

Miami Ls

outer shelf

Pelican Shoal

Holocene coral ridges and sands

Holocene stillstand at 25 m below present sea level

Holocene

lime mud coral ridges and cemented grainstone (~80 ka)

Straits of Florida

Holocene sediments

shelfmargin reef Pleistocene surfaces ~ 1 km

upper-slope terrace (~190 ka) vertical exaggeration 15x

sea level

Miami Ls

TIME

sea level

poorly developed reef or paleoshoreline beach-dune ridge

highstands at ~105, 90, and 80 ka† (Q5 Unit‡, isotope substages 5c, 5b, and 5a§ time)

Hawk Channel coral-ridge growth with sand-filled swales †Chappell, 1974 ‡Perkins, 1977 §Chappell and Shackleton, 1986

Pleistocene surface

shelfmargin reef

upper-slope terrace (~190 ka)

Post-195-ka lowstand and 127-ka transgression phases are similar to those shown in Figure 9. There may be several layers of outer-shelf coral ridges and infilled swales of discrete ages beneath the Holocene layer. sea level Hawk Channel lagoon

~195 ka highstand

initial coral growth on beach-dune ridges inferred paleotopography

NNW

A Nearshore Rock Ledge A prominent, erosional, nearshore rock ledge borders the seaward side of every island in the Florida Keys (Lidz et al. 2003) (Figs. 18.3A, B, 18.13A, B). Its landward edge is the present shoreline. The ledge slopes gently seaward to a depth of ~4 m over a maximum distance of ~2.5 km. Its surface is smooth and marked with potholes. The ledge is generally wider than islands of the Middle and Upper Keys (Fig. 18.13B), consistent with a broad fore reef in front of the emergent reef crests. Its seaward edge is jagged, consistent with a reefal nature and with an irregular fore reef that faced incoming waves (Shinn 1980). The seaward edge forms a vertical scarp whose top is elevated ~30 cm above the Hawk Channel substrate. The raised scarp might be a clue that the ledge is a different type of limestone than that of the channel substrate. Coring the ledge off north Key Largo for installation of water-monitoring wells recovered coral facies like those of the Key Largo Limestone (Shinn et al. 1994). Unfortunately, age of the well coral was not determined. An early concept proposed that the ledge might represent the seaward extent of the Key Largo Limestone and Miami Limestone islands (Lidz et al. 2003). The interpre-

SSE

Figure 18.11. Geomorphogenic model of Pelican Shoal shows a Pleistocene backstepped reef complex at the margin and Holocene backfilled progradation of the outer shelf (top panel; compare with seismic profile in Fig. 18.7A). Note the presence of coral-ridge bands and sediment-filled swales on the outer shelf.

tation was supported by physical characteristics of the ledge, presence of coral facies in the single core, and by consistent presence of the ledge throughout the keys as observed in aerial mosaics. New clear evidence (Lidz et al. 2006) of tidal-bar extent and ledge and channel-substrate contact, enhanced by color hues on the NGDC bathymetric map, supports the interpretation. Coring in the scarp vicinity in Hawk Channel, in transects across and through the ledge base if possible, is needed off islands composed of Key Largo Limestone and Miami Limestone and preferably where the seaward edge of the ledge (the scarp) is most prominent. Examination of the recovered rock and its substrate will prove or disprove the theory. Non-reefal, non-oolitic grainstone and packstone have been cored under Hawk Channel off Key Largo (Shinn et al. 1994; Multer et al. 2002). Core results will also help resolve a long-standing debate on the seaward extent of the Q5 part of the Key Largo Limestone and to a lesser extent the Miami Limestone off the keys (Hoffmeister and Multer 1968; Hoffmeister 1974; Harrison and Coniglio 1985). By definition, the time of formation of an erosional surface is not datable. However, post-125-ka highstand data are a collective proxy: the only possible time the ledge could have been carved into the 125-ka island chain

320 ~ Lidz

Florida Keys

Hawk Channel

Miami Ls

outer shelf Holocene coral ridges and sands

lime mud

Sand Key Reef

Straits of Florida

8.6-7.7 ka† e cen Holo

Holocene sands

shelfmargin reef

coral ridges and cemented grainstone (~80 ka) †Toscano and Lundberg, 1998 ‡Ludwig et al., 1996 §Toscano, 1996; Toscano and Lundberg, 1998, 1999

86.281.6 ka§

primary outlier reef

Pleistocene surfaces

~ 1 km

sea level 8.9-6.7 ka†‡ 83.2-80.9 ka‡ 90.6 ka§ 106.5 ka§

vertical exaggeration 20x

sea level

Miami Ls Hawk Channel

coral-ridge growth with sand-filled swales †Chappell, 1974 ‡Perkins, 1977 §Chappell and Shackleton, 1986

shelfmargin reef outlier reefs on upper-slope terrace (~190 ka)

highstands at ~105, 90, and 80 ka† (Q5 Unit‡, isotope substages 5c, 5b, and 5a§ time)

sea level ~125 ka highstand (Q5 Unit, isotope substage 5e time)

TIME

Miami Ls Hawk Channel depression

coral-ridge growth with sand-filled swales

coral growth

Pleistocene surface

upper-slope terrace (~190 ka)

Post-195-ka lowstand and 127-ka transgression phases are similar to those shown in Figure 9. There may be several layers of outer-shelf coral ridges and infilled swales of different ages beneath the Holocene layer. sea level ~195 ka highstand

Hawk Channel lagoon initial coral growth on beach-dune ridges

inferred paleotopography

NNW

is during the Holocene transgression, a timing that is supported by co-occurrence of the landward edge and shoreline throughout the keys. The highest post-Key Largo Limestone Pleistocene sea level, that of isotope substage 5a, reached its maximum at 9 m below present sea level (Toscano and Lundberg 1999; Lidz 2006) (Table 18.2), which was 5 m lower than the seaward ledge scarp. Ledge width (~2.5 km) and correlation of maximum depth of the ledge surface (~4 m) with a sea-level curve derived from local data (Fig. 18.14A) indicate that island erosion has taken place over a distance of ~2.5 km during the past ~4 thousand years. This equates to ~6.25  m / 100 yr in real time. Given that sea level is rising at a present rate of 0.38 m / 100 yr (Wanless 1989) and is projected to rise as much as 0.88 m in the next 110 yr (Albritton et al. 2001), an erosional interpretation has significant implications for past and future rate of land loss in the densely popu-

SSE

Figure 18.12. Geomorphogenic model of the Sand Key Reef area shows a backstepped and aggradational Pleistocene reef complex at the margin and Holocene backfilled progradation of the outer shelf (top panel). Note bands of outer-shelf coral ridges. Also note that the largest outlier grew upward and landward. Rendition of 3 tracts of outlier reefs is based on seismic data (Fig. 18.6A).

lated Florida Keys. The highest elevation is ~5.5 m (Stanley 1966). Although much of the area is uninhabitable mangrove forest, most inhabitable areas will disappear in a sea-level rise of 1 m (Lidz and Shinn 1991). Under the 0.38-m–rise scenario, the 2-m mark above present sea level would be reached in ~520 years (Fig. 18.14B). Under the projected 0.88-m–rise scenario, the 2-m mark would be reached in ~225 years. The recency, constancy, duration, and distance of erosion under conditions of a rising sea level have immediate relevancy on a real-time, human-interest scale.

Past and Present Platform Progradation Two processes build carbonate platforms seaward: aggradation (slow coral growth) and progradation, generally regarded as the off-bank accumulation of shelf-derived

Coral Reefs of the Florida Keys ~ 321

A

Florida Keys

Florida Bay

scarp sea level lower than elevation of Florida Bay

mangroves scarp nearshore rock ledge

Hawk Channel

Key Largo/Miami Limestone

?

B

grainstone

Florida Bay

Snake Creek

Plan

tatio

n Ke

y

Tavernier Key bank

ck re ro rsho nea ledge

~2.2 km tidal delta

N Hawk Channel

80°31'W

25°N

80°30'W

Figure 18.13. (A) Drawing (from Lidz et al. 2006) shows cross section of nearshore rock ledge as presently inferred. Not to scale. (B) Aerial photo (1975) shows seaward extent of the nearshore rock ledge (short dashed line) in the area of Tavernier Key bank and Plantation Key (Fig. 18.1A) (from Lidz et al. 2003). Note the ledge is much wider than Plantation Key, which may hint at ledge composition. Plantation Key is composed of Key Largo Limestone. Also note the jagged nature of the ledge at its seaward edge, consistent with a reefal structure in plan view. The vertical seaward edge is elevated above Hawk Channel substrate. The landward ledge edge is coincident with the shoreline. Sands (long dashed lines) of Snake Creek tidal delta and sandy lime mud of Tavernier Key bank cover parts of the ledge.

sediments (Hine and Neumann 1977; Hine et al. 1981; Halley et al. 1983; Ginsburg 2001). Classic steeply inclined margins build seaward only by aggradation. The atypical south Florida windward margin accretes through both processes and has done so in alternating areas in the past, as is evident from the data-based models (Figs. 18.9– 18.12) and in the abundance of sand chutes and reentrants in the shelf-edge reef through which sand flows seaward (Figs. 18.3A, 18.6B). Three variations on the standard progradational theme are evident in the Florida reef and sediment record: backstepping reefs, lateral stacking of reefs and infilled troughs, and backfilling of landward troughs (Shinn et al. 1991). The respective stratigraphic results have been termed a backstepped reef-complex

margin, a coalesced reef-complex margin, and a backfilled prograded margin (Lidz 2004). As evidenced by core transects across Looe Key Reef in the Lower Keys, Holocene corals initiated growth on a Pleistocene high composed of massive corals (Shinn et al. 1991) (Fig. 18.1C). The high has the elevation and dimensions of an outlier reef. A sediment-filled trough behind the high and beneath the modern reef was confirmed with drilling and seismic profiles (Lidz et al. 1985). The modern reef has backstepped to a position landward of the crest of the high on which its Holocene counterpart first began growth. Other Holocene outer-shelf reefs have also backstepped in response to rising sea level (Shinn 1980; Shinn et al. 1989; Toscano and Lundberg 1998). Using

322 ~ Lidz

0.23 m/100 yr

1-2 m/100 yr

8.0 m/7 ka = 0.11 m/100 yr avg. for last 8 ka

?

40

Fairbanks (1989) Wanless (1989)

Lighty et al. (1982) Scholl et al. (1969) Fowey Reef coral Acropora palmata

Depth (m) below present sea level

20

Florida Keys data

60

Alligator Reef peat Broad Creek peat

80 Pleistocene

Holocene

the Looe Key Reef stratigraphy as an example, Shinn et al. (1991) proposed that during periods of lowered sea level, berms, dunes, and possibly fringing reefs formed along ancient off-bank shorelines and became subaerially cemented and coated with calcrete. The hardened features provided elevated nuclei for outlier-reef growth as sea level rose. As reefs flourished and grew upward, sediment-lined troughs formed on their landward side. When sea level fell, freshwater diagenesis cemented the reefs and trough sediments, and new berms, dunes, or fringing reefs initiated on a new seaward shoreline and became cemented, creating new topographic highs for outlier-reef growth during the next transgression. Storm waves periodically removed and transported sediment and coral debris into the landward troughs, eventually filling them. The cored coral ridge-and-swale structure near Marker G and the filled-trough structures behind the shelf-edge reef, as modeled from seismic data, support the scenario. Parts of the platform have prograded seaward in a series of sea-level–induced jumps, each of which was as great as the width of the sediment-filled trough and in each of which the outlier reef became the new shelf

Meters above present sea level

1100

1200

0

1000

900

800

700

600

500

0

400

2

300

10

200

12

B

Years into the Future 100

14

Age (ka x 1,000) 8 6 4

1300

A

0.38 m/100 yr 0.04 m/100 yr

6 5 4 3 2 1

IPCC-projected 0.88-m rise in 110 yr Albritton et al. (2001) Present rate of rise 0.38 m/100 yr at Key West

0

Figure 18.14. (A) The sea-level curve for the Florida reef tract is well constrained by local proxy data (in conventional 14 C ages) as modified by Lidz and Shinn (1991) from the curve of Robbin (1984). Data from 8 ka to the present are considered reliable. Rates of sea-level rise at left. Upper part of rise rates shows actual rise measured by tide gauges at Key West (1932-present). (B) Diagram shows the time it will take for sea level to inundate the Florida Keys at its present measured rate of rise (solid diagonal). Dashed line represents predicted rise in mean global sea level from 1990 to 2100 (Albritton et al. 2001) extrapolated to 1 m above present sea level (~110 years) and 2 m above present (~225 years) and beyond. Most dry land in the keys will disappear with a rise in sea level of 1 to 2 m (figure modified from Lidz and Shinn 1991). IPCC = Intergovernmental Panel on Climate Change.

margin. Presence of freestanding outlier reefs off Carysfort Reef and Sand Key Reef implies that new episodes of rapid seaward-stepping progradation will occur when their troughs become filled (Lidz et al. 1991). Stratigraphy resulting from such lateral accretion would represent a coalesced reef-complex margin (Lidz 2004) that would be manifest in cross-sectional seismic profile as parallel coral facies spanning distances of several kilometers. Interpreting this type of margin at an ancient platform might be problematic. Landward- and seaward-dipping bedding would be present, and contacts between coral facies would dip landward. Holocene sediments have filled many areas of the Pleistocene trough behind the present shelf-edge reef, but the stratigraphic result was only formally recognized as margin progradation and termed a backfilled prograded margin when the geomorphogenic models were completed (Lidz 2004). The infill is mostly storm-transported coral rubble derived mainly from the shelf-edge reef with an admixture of seaward-moving sediments from outer-shelf reefs (Shinn et al. 1991). Infilling has been rapid and has occurred solely in the time since the most recent transgression submerged the shelf—during the last 7 to 6 thou-

Coral Reefs of the Florida Keys ~ 323

sand years. Infilling was asymmetric, erasing surface expressions of some (Fig. 18.10), but not all (Fig. 18.9), troughs and leaving large reentrants unaffected (Figs. 18.3A, 18.6B). Infilled areas accreted to or above shelf elevation, thus prograding those parts of the shelf seaward to the new margin, the bare face of the shelf-edge reef.

Marquesas–Quicksands Ridge The Marquesas–Quicksands ridge is a roughly rectangular bedrock structure beneath a sand belt in the Gulf of Mexico known as The Quicksands (Shinn et al. 1990) (Fig. 18.1A). Relatively deep (≥20 m) waters bound the ridge on 3 sides, marked by the 10-m depth contour: the Tortugas or Key West Shrimping Grounds to the north, an unnamed channel to the west, and a back-reef lagoon to the south—the westward extension of the Hawk Channel bedrock depression (Fig. 18.1C). The floor of Boca Grande Channel, a shallow (4.5 to 6 m) north-trending bedrock depression east of the ridge, is separated from similar ridge depths by high bedrock beneath the Marquesas Keys (Shinn et al. 1990; Lidz et al. 2003). Strong reversing tidal currents between the Gulf of Mexico and Straits of Florida keep Boca Grande Channel free of sediment. The broad (~28 by 47  km) ridge is the submerged extension of the arcuate surface that forms the Florida Keys to the east and extends beyond the Dry Tortugas to the west. Capped by calcrete, the bedrock ranges from 0 to 12 m below sea level but is typically ~6 to 8 m deep with numerous isolated depressions and solution holes (Lidz et al. 2003). Oolite has been recovered in core borings around the Marquesas Keys and is believed to extend as far westward as Halfmoon Shoal (Shinn et al. 1990). Beneath and south of the islands, the oolite is non-bedded and highly burrowed. North of the islands, it is distinctly cross-bedded with no burrows. Cross-bedding and a northward dip are characteristic of a beach setting that existed at a lower sea level (Shinn et al. 1989). Ellis Rock and New Ground Shoal on the north side of the ridge (Fig. 18.1C) are parts of a narrow line of Holocene reefs with intervening carbonate muds and silts as thick as 7 m. From rotary rock cores drilled at both sites, the intermittent shoals are known to consist of massive relict head corals ~7.6 m thick that overlie corals typical of the Key Largo Limestone (Shinn et al. 1990). This line of reefs would have been subject to the strong currents and the damaging effects of cold water temperatures in the Gulf of Mexico (Mayer 1914; Roberts et al. 1982). North

of the reef line, water depth drops from about 6 m at New Ground Shoal to an average of 23 m, and the sediments change from coral framework to the lime mud and silt of the Tortugas Shrimping Grounds. Halfmoon Shoal and Rebecca Shoal are bedrock highs located on the respective southwestern end of the ridge and northwestern side of the unnamed channel (Fig. 18.1A). Both shoals may represent old patch reefs, but their origins have not been studied. The Quicksands, a large volume of non-oolitic carbonate sand, migrate westward across the ridge despite north–south movement of the strong currents. The sand consists primarily of fragmented Halimeda plates (Hudson 1985). On average 1 to 2 m thick and molded by the currents, the sand-belt surface is ornamented with large (5 m high) north-trending tidal bars topped by smaller (1 m high) west-trending sand waves (Shinn et al. 1990). Some sand-wave crests are awash at low tide. The lowamplitude sand waves overlie bedrock near the Marquesas Keys. The high-amplitude waves at the west end of the ridge overlie westward-dipping sand beds as thick as 12 m. Westward and off-bank sediment transport in the Lower Keys and a sediment-free Boca Grande Channel signify that the entire sand body probably was, and still is, being generated on the ridge.

Future Margin Evolution: Increasing Intricacy Perkins (1977) regarded the overall Pleistocene record of south Florida as having a tendency toward leveling of pre-existing topography. Composite results of the database synthesis show his analysis cannot apply to late Quaternary shelf-edge stratigraphy (Lidz et  al. 2003; Lidz 2004). Where reef complexes did not build up at the margin, reentrants have remained much as they were when formed, resulting in a “notched” margin. Alternating parts of the margin have prograded through a combined seaward stacking of reefs and infilled troughs and Holocene backfilling of Pleistocene troughs. The presence of outlier reefs off the shelf-edge reef, but not off large reentrants such as at Alligator Reef and Tennessee Reef (Lidz et al. 2003) (Fig. 18.1C), implies increasing intricacy or irregularity of the shelf edge through time. Where outlier reefs and back-reef troughs exist seaward of the present shelf-edge reef, those parts of the shelf and margin will build seaward primarily in sea-level–controlled jumps through infilling of the troughs. Under the right conditions, as is evident in the Holocene record, infilling requires only a few thousand years. These later-

324 ~ Lidz

ally accreting areas will form promontories of coalesced reef complexes between the reentrants (Lidz 2004). The recurring theme is the replication of antecedent facies and geomorphic landforms, and thus repetition of stratigraphic shelf-edge asymmetry. Where outlier reefs are absent, the margin will accrete seaward more slowly through sediment accrual in the reentrants. Unfilled terrace areas may eventually become new reentrants if sea level falls and the sands become cemented. If sea level were to remain high and stable over hundreds of thousands of years, all accommodation space on the shelf would theoretically fill to the sea surface (Tucker 1985). The exception would be if the depositional environment were able to export more sediment than it produced or received from elsewhere. Physically and biologically predisposed, the late Quaternary depositional system along the shallow Florida shelf has replicated the structure and form of antecedent shelf-edge morphologies that are distinctly different from those that would be generated at a classic steeply inclined windward margin. Morphogenic definition of the seismic stratigraphic sequences provides perspectives on characterizing reef-complex buildup in general. The perspectives serve as a guide and stepping-stone for deciphering accretionary processes at other reef-rimmed carbonate platforms, whether windward or leeward, modern or ancient.

Conclusions Thirteen transgressive–regressive cycles controlled late Quaternary deposition at the south Florida windward margin between ~350  ka and today. Seven successive coral reef complexes evolved on top of 2 non-coralline marine sequences (Q1 and Q2 units) and produced 6 major coral structures: the inland Key Largo Limestone (and associated ooid tidal-bar Miami Limestone) of the Florida Keys, a shelf-edge reef, and 4 tracts of outlier reefs on an upper-slope terrace. The stratigraphic Q3 and Q4 units are mid-Pleistocene. The shelf-edge and outlier reefs, dated between ~106.5 and ~80 ka, are inferred to contain corals at depth that would date to the 127-ka transgression. Corals in the Q5 Key Largo Limestone date to marine-isotope substage 5e. Offshore reefs date to substages 5c, 5b, and 5a and to Stage 1. Sea level remained below the elevation of the Florida shelf between substage-5a time (youngest coral date ~77.8 ka) and the Holocene (oldest coral date ~9.6 ka), during which time

mangrove peat and calcrete, now offshore, accumulated. Actual duration of this exposure event was ~115 thousand years, as evidenced by the highest post-125-ka transgression reaching a maximum elevation of only 9 m below present sea level. Corals recruited to 4 types of calcrete-coated nucleate substrates: antecedent reefs and beach-dune ridges along the outer shelf and upper-slope terrace, and non-coralline topographic-trough margins and the oolitic and reefal nearshore rock ledge on the inner shelf. Pleistocene reefs grew on the first 2. Holocene reefs grew on all 4. Beachdune ridges are inferred to underlie all linear reefs. On the basis of geomorphic elements, 2 basic coral reef structures formed off the Florida Keys. A narrow bank-top ridge-and-swale architecture built the outer shelf upward and seaward. A broad shelf-edge reef and infilled-trough architecture built alternate parts of the margin upward, landward, and seaward. The processes were both aggradational and progradational. Massive head corals dominated the Pleistocene. Framework-building branching acroporids dominated the Holocene. In comparison, coral growth on the Marquesas–Quicksands ridge to the west was largely sparse. Alternating non-reentrant shelf-edge areas have been prograded seaward through systematic vertical and lateral stacking of reefs and infilled back-reef troughs. Replication of antecedent facies and geomorphic landforms has increased margin asymmetry through time. Geologic imprint of the Holocene transgression is twofold. Parts of the shelf surface have prograded or stepped seaward in less than 7 to 6 thousand years through infilling of Pleistocene back-reef troughs. The seaward sides of the Florida Keys have eroded landward at a rate of ~6.25 m / 100 yr over the past ~4 thousand years, forming a 2.5-km–wide landward-sloping nearshore rock ledge. Implications of the persistency of this erosion under a present rise in sea level and a projected increase in rate of rise have real-time relevancy for residents of the keys. The Quaternary windward margin of Florida has a spatial geomorphic complexity unlike other windward margins in the Caribbean and Bahamas. Present discontinuity of shelf-edge morphologies and future differential accretion of reefs and sediments will increase lateral shelf-edge asymmetry through time. Yet a regular, consistent stratigraphic simplicity in mode of evolution is evident in geomorphogenic models. Paleotopography and fluctuating sea level controlled the processes and resultant geomorphic architectures. The cause, at least in part, of the uniqueness of this margin among windward margins

Coral Reefs of the Florida Keys ~ 325

is probably its curvature and low-gradient slope. Most carbonate platforms consist of many thousands of meters of accreted carbonates, or volcanoes, both of which have precipitous slopes.

Acknowledgments Gratitude is extended to Jim Pitts of the Department of Transportation in Tallahassee, Florida, who kindly supplied the 1975 aerial photographs, and to Joan Rikon and Patricia Kibler of NOAA National Geodetic Survey in Silver Spring, Maryland, who provided the 1991 photographs. Both photo sets encompass the area of the Florida Keys and shelf to a depth of ~40 m and extend from off central Key Largo to the west end of the Marquesas– Quicksands ridge in the Gulf of Mexico. Both sets were used to construct the habitat map of Lidz et al. (2006) on which certain discussions in this chapter are based. In some cases, the photos provided the sole evidence for a particular part of the geologic record, such as the presence of 4 tracts of outlier reefs shown in Figure 18.6B that would otherwise not have been known from other types of data. The photos in Figures 18.3A, B, 18.6B, and 18.13B are from the 1975 set. The author thanks Brian Bossak and John Brock for constructive reviews of the manuscript. The USGS Coastal and Marine Geology Program funded the data-synthesis project from which this summary paper is derived. The 1997 seismic geophysical data were acquired under Florida Keys National Marine Sanctuary Permit FKNMS-27–97. All products using or referencing any portion of the 1997 data are attributable to and should acknowledge that permit number.

References Albritton, D. L., M. R. Allen, A. P. M. Baede, and 56 others. 2001. Summary for Policymakers. Pp. 1–20 in J. T. Houghton, Y. Ding, D. J. Griggs, M. Noguer, P. J. van der Linden, X. Dai, K. Maskell, and C. A. Johnson (eds.), Climate Change 2001: The Scientific Basis, Contribution of Working Group I to the Third Assessment Report of the Intergovernmental Panel on Climate Change. New York: Cambridge University Press. Aronson, R. B., and W. F. Precht. 1997. Stasis, biological disturbance, and community structure of a Holocene coral reef. Paleobiology 23:326–46.

Ball, M. M. 1967. Carbonate sand bodies of Florida and the Bahamas. Journal of Sedimentary Petrology 37:556–91. Ball, M. M., E. A. Shinn, and K. W. Stockman. 1967. The geologic effects of Hurricane Donna in South Florida. Journal of Geology 75:583–97. Bard, E., B. Hamelin, R. G. Fairbanks, A. Zindler, G. Mathieu, and M. Arnold. 1990. U / Th and 14C ages of corals from Barbados and their use for calibrating the 14C time scale beyond 9000 years B.P. Nuclear Instruments and Methods in Physics Research B52:461–68. Bloom, A. L., W. S. Broecker, J. M. A. Chappell, R. K. Matthews, and K. J. Mesolella. 1974. Quaternary sea level fluctuations on a tectonic coast: new 230Th / 234U dates from the Huon Peninsula, New Guinea. Quaternary Research 4:185–205. Buddemeier, R. W., and S. V. Smith. 1988. Coral reef growth in an era of rapidly rising sea level: predictions and suggestions for long-term research. Coral Reefs 7:51–56. Causey, B. 2002. The role of the Florida Keys National Marine Sanctuary in the South Florida Ecosystem Restoration Initiative. Pp. 883–94 in J. W. Porter and K. G. Porter (eds.), The Everglades, Florida Bay and Coral Reefs of the Florida Keys: An Ecosystem Sourcebook. Boca Raton, Fla.: CRC Press. Chappell, J. 1974. Geology of coral terraces, Huon Peninsula, New Guinea: a study of Quaternary tectonic movements and sea-level changes. Geological Society of America Bulletin 85:553–70. Chappell, J., and N. J. Shackleton. 1986. Oxygen isotopes and sea level. Nature 324:137–40. D’Almeida, G. A. 1986. A model for Saharan dust transport. Journal of Climatology and Applied Meteorology 24:903–16. Divins, D. L. 2003. Coastal Relief Model. Boulder, Colo.: National Geophysical Data Center. http: // www.ngdc.noaa .gov / mgg / coastal / coastal.html Dustan, P. 1985. Community structure of reef-building corals in the Florida Keys: Carysfort Reef, Key Largo and Long Key Reef, Dry Tortugas. Atoll Research Bulletin 288:1–27. Enos, P. 1977. Holocene sediment accumulations of the South Florida shelf margin. Pp. 1–130 in P. Enos and R. D. Perkins (eds.), Quaternary Sedimentation in South Florida. Memoir 147. Boulder, Colo.: Geological Society of America. Enos, P., and R. D. Perkins (eds.). 1977. Quaternary Sedimentation in South Florida. Memoir 147. Boulder, Colo.: Geological Society of America. 198 pp. EPA (Environmental Protection Agency). 1983. Florida Keys seagrass and coral reef inventory, vol. 2, Sombrero Key to Marquesas Rock, February 1982. Report to U.S. Fish and Wildlife Service and Minerals Management Service (Contract No. 68–03–3049). Las Vegas, Nev.: EPA Environmental Monitoring Systems Laboratory. 104 pp.

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Lidz, B. H., and E. A. Shinn. 1991. Paleoshorelines, reefs, and a rising sea: South Florida, U.S.A. Journal of Coastal Research 7:203–29. Lidz, B. H., E. A. Shinn, M. E. Hansen, R. B. Halley, M. W. Harris, S. D. Locker, and A. C. Hine. 1997a. Maps showing sedimentary and biological environments, depth to Pleistocene bedrock, and Holocene sediment and reef thickness from Molasses Reef to Elbow Reef, Key Largo, South Florida. Miscellaneous Investigations Series Map #I-2505. Washington, D.C.: U.S. Geological Survey. 3 sheets, scale 1:24,000. Lidz, B. H., E. A. Shinn, A. C. Hine, and S. D. Locker. 1997b. Contrasts within an outlier-reef system: evidence for differential Quaternary evolution, South Florida windward margin, U.S.A. Journal of Coastal Research 13:711–31. Lighty, R. G., I. G. Macintyre, and R. Stuckenrath. 1978. Submerged early Holocene barrier reef, southeast Florida shelf. Nature 276:59–60. ———. 1982. Acropora palmata reef framework: a reliable indicator of sea level in the western Atlantic for the past 10,000 years. Coral Reefs 1:125–30. Locker, S. D., and A. C. Hine. 1995. Late Quaternary Sequence stratigraphy, south Florida margin. Pp. 319–27 in Proceedings of the 27th Annual Ocean Technology Conference, Volume 1, Geology, Earth Sciences & Environmental Factors. Dallas, Tex. Locker, S. D., A. C. Hine, L. P. Tedesco, and E. A. Shinn. 1996. Magnitude and timing of episodic sea-level rise during the last deglaciation. Geology 24:827–30. Ludwig, K. R., D. R. Muhs, K. R. Simmons, R. B. Halley, and E. A. Shinn. 1996. Sea-level records at ~80 ka from tectonically stable platforms: Florida and Bermuda. Geology 24(3):211–14. Lundberg, J., and D. C. Ford. 1994. Late Pleistocene sea-level change in the Bahamas from mass spectrometric U-series dating of submerged speleothem. Quaternary Science Reviews 13:1–14. Mallinson, D., A. Hine, P. Hallock, S. Locker, E. Shinn, D. Naar, B. Donahue, and D. Weaver. 2003. Development of small carbonate banks on the South Florida platform margin: response to sea level and climate change. Marine Geology 199:45–63. Marszalek, D. S. 1977. Florida reef tract marine habitats and ecosystems. Maps published in cooperation with State of Florida Department of Natural Resources; U.S. Department of the Interior Bureau of Land Management, New Orleans Outer Continental Shelf Office; and University of Miami Rosenstiel School of Marine and Atmospheric Science. 9 sheets, scale 1:30,000.

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Mayer, A. G. 1914. The effects of temperature on tropical marine animals. Carnegie Institute of Washington Publication 183:6:1–24. McKee, E. D., and W. C. Ward. 1983. Eolian environment. Pp. 131–70 in P. A. Scholle, D. G. Bebout, and C. H. Moore (eds.), Carbonate Depositional Environments. Memoir 33. Tulsa, Okla.: American Association of Petroleum Geologists. Milliman, J. D., and K. O. Emery. 1968. Sea levels during the past 35,000 years. Science 162:1121–23. Mitterer, R. M. 1974. Pleistocene stratigraphy in southern Florida based on amino acid diagenesis in fossil Mercenaria. Geology 2:425–28. Muhs, D. R. 2002. Evidence for the timing and duration of the last interglacial period from high-precision uranium-series ages of corals on tectonically stable coastlines. Quaternary Research 58:36–40. Muhs, D. R., C. A. Bush, and K. C. Stewart. 1990. Geochemical evidence of Saharan dust parent material for soils developed on Quaternary limestones of Caribbean and Western Atlantic islands. Quaternary Research 33:157–77. Muhs, D. R., B. J. Szabo, L. McCartan, P. B. Maat, C. A. Bush, and R. B. Halley. 1992. Uranium-series age estimates of corals from Quaternary marine sediments of southern Florida. Pp. 41–49 in T. M. Scott and W. D. Allmon (eds.), Plio-Pleistocene Stratigraphy and Paleontology of Southern Florida. Special Publication no. 36. Tallahassee, Fla.: Florida Geological Survey. Muhs, D. R., J. F. Wehmiller, K. R. Simmons, and L. L. York. 2004. Quaternary sea-level history of the United States. Pp. 147–84 in A. R. Gillespie, S. C. Porter, and B. F. Atwater (ed.), The Quaternary Period in the United States, Developments in Quaternary Science, vol. 1. Maryland Heights, Mo.: Elsevier. doi:10.1016 / S1571–0866(03)01008-X. Multer, H. G. 1977. Field Guide to Some Carbonate Rock Environments, Florida Keys and Western Bahamas, 6th ed. Dubuque, Iowa: Kendall Hunt Publishing. 425 pp. ———. 1996. Seakeys habitat guide to the Florida Keys National Marine Sanctuary: past, present, future (double-sided colored descriptive and interpretive map, 1 sheet, scale 1:400,000). Copies available from: The Florida Keys National Marine Sanctuary, P.O. Box 500368, Marathon, Fla. 33050; e-mail: www.fknms.nos.noaa.gov. Multer, H. G., E. Gischler, J. Lundberg, K. R. Simmons, and E. A. Shinn. 2002. Key Largo Limestone revisited: Pleistocene shelf-edge facies, Florida Keys, USA. Facies 46:229–72. Multer, H. G., and J. E. Hoffmeister. 1968. Subaerial laminated crusts of the Florida Keys. Geological Society of America Bulletin 79:183–92.

Neumann, A. C., and I. G. Macintyre. 1985. Response to sea level rise: keep-up, catch-up, or give-up. Proceedings of the Fifth International Coral Reef Congress 3:105–10. Perkins, R. D. 1977. Depositional framework of Pleistocene rocks in South Florida. Pp. 131–98 in P. Enos and R. D. Perkins (eds.), Quaternary Sedimentation in South Florida. Memoir 147. Boulder, Colo.: Geological Society of America. Pratson, L., D. Divins, T. Butler, D. Metzger, M. Steele, G. Sharman, T. Berggren, T. Holcombe, and R. Ramos. 1999. Development of an elevation database for the U.S. Coastal Zone. Survey Land Information Systems 59:3–13. Robbin, D. M. 1981. Subaerial CaCO3 crust: a tool for timing reef initiation and defining sea level changes. Proceedings of the Fourth International Coral Reef Congress 1:575–79. ———. 1984. A new Holocene sea level curve for the upper Florida Keys and Florida reef tract. Pp. 437–58 in P. J. Gleason (ed.), Environments of South Florida: Present and Past. Memoir 2. Miami, Fla.: Miami Geological Society. Roberts, H. H., L. J. Rouse Jr., N. D. Walker, and J. H. Hudson. 1982. Cold-water stress in Florida Bay and northern Bahamas; a product of winter cold-air outbreaks. Journal of Sedimentary Petrology 52:145–55. Sanford, S. 1909. The topography and geology of southern Florida. Florida Geological Survey Annual Report 2: 175–231. Scholl, D. W., F. C. Craighead Sr., and M. Stuiver. 1969. Florida submergence curve revised: its relation to coastal sedimentation rates. Science 163(3867):562–64. Schroeder, R. A., and J. L. Bada. 1976. A review of the geochemical applications of the amino acid racemization reaction. Earth-Science Reviews 12:347–91. Shinn, E. A. 1963. Spur and groove formation on the Florida reef tract. Journal of Sedimentary Petrology 33:291–303. ———. 1980. Geologic history of Grecian Rocks, Key Largo Coral Reef Marine Sanctuary. Bulletin of Marine Science 30:646–56. ———. 1988. The geology of the Florida Keys. Oceanus 31:46–53. Shinn, E. A., J. H. Hudson, R. B. Halley, and B. H. Lidz. 1977. Topographic control and accumulation rate of some Holocene coral reefs, South Florida and Dry Tortugas. Proceedings of the 3rd International Coral Reef Symposium 2:1–7. Shinn, E. A., B. H. Lidz, and A. C. Hine. 1991. Coastal evolution and sea-level history: Florida Keys. Pp. 237–39 in Coastal Depositional Systems in the Gulf of Mexico, Quaternary Framework and Environmental Issues: 12th Annual GCSSEPM Foundation Research Conference. Houston, Tex.: GCSSEPM Foundation.

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␥19

Florida Middle Ground Reef Complex Gregg R. Brooks and David Mallinson

The Florida Middle Ground (FMG) reef complex, located on the west Florida continental margin (see Hine and Locker in chapter 7 in this volume) in the northeastern Gulf of Mexico, is unusual in that it occupies a relatively high-latitude (>28°N), mid-shelf setting (Fig. 19.1). As one of the most productive fishing grounds in the Gulf of Mexico since the 1880s, most early research on the FMG was biological. Investigations of the physical environment began in the 1950s, but most of the initial work was descriptive and focused on reef physiography primarily for navigation for the fishing fleet. Geological studies began in the 1970s and dealt principally with sediments and sedimentary processes. The most recent investigations, beginning in the 1990s, have utilized the vast improvements in available technology to fine-tune and improve upon previous investigations. Current investigations focus on timing and controls of FMG development and how they relate to paleoclimates and Gulf of Mexico circulation patterns. This chapter provides a broad overview of the FMG. The key geologic findings over the past 50 years are summarized and gaps in our knowledge are identified. The chapter begins with an overview of the physical setting of the FMG followed by a discussion of the regional geology and structural framework. We then focus on the FMG as a biological and geological system and discuss controls on development throughout the recent geologic past. Since reef systems are both biological and geological features, both biological and geological processes will be reviewed.

Setting The FMG covers over 900 km2 of the west Florida continental shelf (28°10′N–28°40′N; 84°5′W–84°25′W) and water depths vary between 24 and 45 m (Fig. 19.2). The highly variable morphology of high- and low-relief live bottoms occur along a 60 km long by 15 km wide area that trends north-northwest, parallel to the platform margin. Individual ridges are approximately 12–15 m in height, as much as 2 to 3 km wide, and are separated by a broad, relatively flat valley up to 7 km in width. The southern portion of the FMG is characterized by low relief, with depths ranging from 34 to 45 m. The northern portion of the FMG is characterized by greater relief, with depths ranging from 24 to 45 m. The carbonate banks are surrounded by rock rubble and sediment aprons. Diver reconnaissance revealed highly bioeroded margins with abundant cryptic environments, including swim-throughs, overhangs, and shallow caverns. The geologic setting, that of a mid- to outer-shelf carbonate buildup in a ramp, is similar to the reefs of Campeche Bank on Mexico’s Yucatan Peninsula (Logan et al. 1969; Hine and Mullins 1983).

Regional Geology The FMG is located on the western flank of the Florida carbonate platform (see Scott in chapter 2 in this volume),

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Figure 19.1. Map of the eastern Gulf of Mexico showing the location of the Florida Middle Ground on the west Florida continental shelf.

which was once part of the Florida–Bahamas “gigaplatform.” During the Jurassic, this enormous feature extended from the Gulf of Mexico to the Grand Banks and accumulated up to 12 km of shallow-water carbonate sediments (Poag 1991). Basement rocks beneath the west Florida shelf consist of a thick, transitional crust with up to 5 km of relief, forming a series of arches and embayments (Sawyer et al. 1991; Randazzo 1997). One of these structures, the Middle Ground Arch, provides the foundation for the FMG reef complex and may have supported reefs 100 million years ago (Winston 1969; Brooks 1974). The Mesozoic to Cenozoic carbonate cover overlying these basement rocks reaches 2–7 km in thickness, with little to no bathymetric expression of the Middle Ground Arch on the modern shelf surface. The west Florida shelf surface in this region consists of flat to gently rolling topography described as a drowned karst plain with solution basins and sinkholes (Brooks 1974). Sinkholes were likely formed during sea-level lowstands and subsequently modified by infilling with coastal sediments as sea level transgressed the shelf surface (Price 1954). Soft sediment cover in the FMG area ranges from 0 to 10 m thick. For a more detailed discussion of the west Florida continental shelf see Hine and Locker in chapter 7 in this volume.

Physical Oceanography The FMG occupies a complex zone of mixing between the Caribbean-derived Loop Current, the West Florida Estua-

Florida Middle Ground Reef Complex ~ 333

rine Gyre, and Florida Bay waters (Austin and Jones 1974) (Fig. 19.3). The water is generally clear, of stable salinity (36 ± 0.5‰), and is classified as north tropical (Austin and Jones 1974; Cheney and Dyer 1974). Average annual bottom temperatures fluctuate between 16 °C and 20 °C with lowest temperatures in January and February and the highest temperatures in September and October. During the summer, current-meter data indicate a 2-layered flow with the thermocline acting as the shear zone (Gaul 1967; Mofjeld 1974). As spring approaches, the warmer temperatures cause a weak, vertical thermal gradient to form, which becomes stronger as the season progresses. Surface temperatures average about 23  °C whereas bottom temperatures average 19  °C to 20  °C. Water below the thermocline is believed to be associated with the Loop Current (Gaul 1967). The surface water, however, is derived from the West Florida Estuarine Gyre and is periodically replaced by eddies that spin off from the northern boundary of the Loop Current (Austin and Jones 1974). Direct observation by divers, as well as tidally periodic displacement of the thermocline, suggest internal waves also are present during the summer (Gaul 1967; Mofjeld 1974).

N

During winter, circulation is not as dependent upon the driving forces of the Loop Current. At this time, the local wind conditions and, to some extent, tidal forces are the primary controlling factors (Austin and Jones 1974). Bottom currents fluctuate widely, but velocities of over 100 cm / s have been observed near the FMG ridge tops (Back 1972). Highest values are reported during summer and early autumn and may be associated with the strong thermocline. Strong bottom currents at depths greater than 34 m are also indicated by the occurrence of scour channels, ripple marks, and sand waves (Back 1972; Mallinson et al. 2006). Upwelling is a common occurrence on the outer margins of the Loop Current where seasonal high-chlorophyll plumes produce mesotrophic to eutrophic conditions. Upwelling, along with the complex circulation patterns, provides a mechanism for mixing of nutrient-rich waters and has been cited as a major factor responsible for high productivity on the FMG (Williams 1972; Austin and Jones 1974). Additionally, these nutrient resources are important in limiting coral-reef development in the FMG area (Margalef 1968; Kinsey and Davies 1979; Wiebe et al. 1981; Hallock and Schlager 1986; Birkeland 1987). Hallock (1988) proposed that excess nutrients might stimulate bioerosion and suppress reef development on the west Florida shelf. At present, the FMG area is affected by periodic upwelling (Austin and Jones 1974) and seasonal chlorophyll plumes indicating relatively high-nutrient water conditions (Gilbes et al. 1996). These plumes have been recognized each spring using Coastal Zone Color Scanner imagery. Plumes begin north of the FMG area near Apalachicola Bay and migrate south-southeast directly over the FMG. Mechanisms proposed for the origin of the plume include interactions of the Loop Current with the platform margin producing upwelling, and entrainment of high-nutrient water masses from fluvial discharge. Present environmental conditions support zooxanthellate coral growth (Grimm and Hopkins 1977) but are not sufficiently favorable to support continued reef growth in the FMG area.

Overview of FMG Investigations: 1950s to Present Figure 19.3. Map showing the Florida Middle Ground at the intersection of the Loop Current, Florida Bay waters, and the West Florida Estuarine Gyre (modified from Austin and Jones 1974). Stippled rectangle represents area of Florida Middle Grounds.

The FMG has been known by fishermen as far back as the 1880s as one of the most productive fishing grounds in the Gulf of Mexico. A reconnaissance survey was made by the coast and geodetic survey in 1914 in response to

334 ~ Brooks and Mallinson

fishermen’s reports of shoal depths, but it was not until the 1950s that investigations began on the physical environment (primarily geology and physical oceanography). The first systematic mapping of the FMG was made by Jordan (1952). On the basis of numerous fathometer profiles, he produced a bathymetric map contoured at 1-fathom intervals and described what he called a distinctive reef formation having the appearance of an old river delta. It was described as having a 3-mile-wide “river bed” along a north–south axis, flanked by embankments up to 50 feet high. The “mouth” of the riverbed occupied the seaward end, which was divided by an extensive shoal or reef area with small reefs dotting the passes (Jordan 1952). Samples obtained from the tops of embankments and shoals consisted of a large amount of coral debris. The irregular coral bottom provided an ideal feeding ground for fish, explaining why the FMG was the foundation for the Pensacola and Tampa fishing fleet for red snapper and grouper during this period (Jordan 1952). The foundation for the FMG reefs has always been a mystery, but it was originally thought to be the irregular, karst surface of the Florida carbonate plateau (Price 1954). It was finally confirmed that relief features were due to reef growth after numerous samples of calcareous algae were collected from the FMG by Gould and Stewart (1956). Direct observations started in the 1960s when Brooks (1962) studied the area using scuba and found living scleractinians, hydrocorals, alcyonarians, coralline algae, and fish similar to those on Florida Keys reefs. These organisms had not previously been recorded that far north or at those depths. The foundation for the modern reef was likely an older reef that may have developed and flourished during the lower sea levels of the last interglacial (Brooks 1962). In the late 1960s, large-scale interdisciplinary studies were initiated on the FMG. Austin (1970) reported that living and dead corals on the FMG are typical of deep Caribbean reefs, and the invertebrates and fish along with the corals reflect an impoverished coral-reef fauna. A study of the physical oceanography revealed that the FMG occupies a region of climatological convergence, and therefore, a transition zone among the Gulf Loop Current, the West Florida Estuarine Gyre, and Florida Bay waters (Austin 1970) (Fig. 19.3). Geologic investigations began in earnest in the 1970s. Back (1972) investigated the recent depositional environment of the FMG based upon bathymetry, sediment texture and composition, the benthos, bottom-current data,

and general ecological observations. He concluded that the predominant sediment sources for the FMG area are the reef ridges and knolls where the vigorous growth of carbonate-secreting organisms is greatest. The important biogenic sediment contributors include mollusks, Lithothamnion algae, barnacles, foraminifera, madreporian corals, and bryozoans. Carbonate sediments are winnowed off the reef ridges and knolls and transported to the flat plain where they are deposited and mixed with fine quartz sands. The reef ridges and knolls typically consist of a dense, 1-m-high Millepora zone near the reef edges, which to some degree act as a rim inhibiting the removal of reef sediments. Strong bottom currents were recorded and are considered responsible for winnowing material from the reef knolls and pinnacles. These bottom currents are also likely responsible for producing a gently sloping, 2-m-deep scour channel surrounding many of the banks, and large bedform fields in the northern FMG area (Mallinson et al. 2006). Back (1972) concluded that the shoal portions of the FMG may be considered coral reefs on the basis of diversity and abundance of organisms, but there is no evidence of significant modern reef growth. Analysis of a rock core, 2.1 m into a reef crest, produced an age of approximately 6.5 kyr for FMG reefs (Back 1972). This yielded an average reef growth of approximately 27 cm / kyr, which is consistent with other similar reef systems (Back 1972). In the late 1970s, a large-scale, multidisciplinary program was initiated on the FMG as part of the Gulf of Mexico Topographic Features Study. Geologic investigations included geophysical mapping, sediment texture and composition, foraminiferal analysis, and direct observation using a submersible. Geophysical surveys yielded 3 seismic units in the study area. The basal unit was characterized by a very irregular, highly reflective surface representing the acoustic basement. It was interpreted as the Miocene karst surface underlying the surficial unconsolidated sediment veneer throughout the west Florida shelf. The middle unit, bounded on the top by a horizontal surface, was interpreted to reflect the Holocene transgressive sand sheet. This unit downlaps and pinches out against the underlying surface north-northwest of the study area (Hilde et al. 1981). The surficial unit thickens in the middle of the central valley, as well as on the flanks where it onlaps the reef ridges. It was interpreted as sheet drape, with sediments likely originating from the adjacent reef ridges (Hilde et al. 1981). Sediment samples, collected at 4 stations to establish the depositional framework for biological investigations, consisted of carbonate

Florida Middle Ground Reef Complex ~ 335

sands dominated by molluscan shell fragments similar to those of the surrounding shelf (Doyle et al. 1980). Either the sediments have been masked by shelf sediments swept into the study area or little or no carbonate production is occurring on reef ridges. Direct observations from a submersible support the latter conclusion because reef growth has been primarily vertical with little horizontal spreading (Doyle et al. 1980). Geological investigations continued into the early 1980s when Brooks (1981) found that, with the exception of an abundance of barnacle fragments, FMG sediments were more like those of the surrounding shelf environment than a typical coral reef. From the surficial sediments, Brooks (1981) concluded that the FMG was in a transitional phase between a reef and open carbonate shelf environment. In the 1990s investigations began to focus on factors controlling the geologic development of the FMG. Integrating sedimentology with high-resolution seismicreflection data and direct observations from a submersible, Brooks and Doyle (1991) concluded that the recent geologic development of the FMG has been driven by high-frequency sea-level fluctuations. The thick surface sediments associated with the FMG, relative to the surrounding shelf, indicate that reef ridges act to trap sediments much like shelf-edge reefs trap sediments on rimmed carbonate platforms. Reef growth and sediment accumulation have not kept pace with sea-level rise, which is drowning FMG reefs along with the remainder of the west Florida carbonate margin (Brooks and Doyle 1991). The most recent investigations on the FMG, beginning in the late 1990s, have utilized the vast improvements in available technology. More detailed detection and mapping of benthic environments have fine-tuned and improved upon previous investigations, and the acquisition of high-resolution seismic data has refined the understanding of the stratigraphic framework. These data are being used to help determine the foundation and timing of original reef formation. On the basis of multibeam side-scan sonar, high-resolution sub-bottom profiling data, and diver and ROV (remotely operated vehicle) observations, Mallinson (2000) and Mallinson et al. (2006) identified bioeroded margins on the FMG with abundant cryptic environments including arches, overhangs, and shallow caverns. The benthic biota was dominated by Millepora alcicornis, Dichocoenia stokesii, octocorals, Muricea spp., Madracis decactis, Xestospongia muta, and tube sponges. They also reported a complex

geologic framework of relict Pleistocene limestone mantled with Holocene corals and sediment. The high-relief carbonate banks were likely developed on low-relief environments and represent zones of maximum early to midHolocene reef recruitment and growth. The present lowrelief environments are likely Pleistocene shallow-water limestones that were intermittently exposed to subaerial weathering processes during Pleistocene lowstands and subsequently mantled with Holocene corals (Mallinson et al. 2006). Seismic and side-scan sonar data (Figs. 19.4, 19.5) revealed the Miocene basement reflection exhibiting 5  ms in karstic relief (Mallinson et  al. 2006), supporting previous interpretations by Hilde et al. (1981). Several seaward-dipping Plio-Pleistocene sequences were bounded at the top by an angular unconformity. Reef location appears to be related to the sub-crop pattern of these units along the angular unconformity (Fig. 19.4). The reefs extend upward from shallow, subsurface reflections and assume several morphologies, including steep-sided, flat-topped banks and pinnacles. The presence of the FMG may indicate the existence of warmer, more oligotrophic conditions in the past, either during the late Pleistocene or mid-Holocene (Hilde et al. 1981). Mallinson (2000) defined acoustic-backscatter facies and related them to bottom type (Fig. 19.6). Live-bottom communities are recruited upon hard-bottom highs, evidence that bottom type and bathymetry are correlated. Live bottoms and hard-bottom highs are characterized by a high acoustic-backscatter signature with a patchy, crenulated texture. Diver and underwater camera reconnaissance revealed that live-bottom communities are dominated by various sponges, gorgonian corals, hydrozoans, small (8 km) margin-parallel banks 2 to 3 km in width, (2) isolated flat-topped banks with no particular orientation or maximum dimension, and (3)  isolated patch reefs generally 63 μm residues were dry sieved and separated into 63 μm150 μm and >150 μm size fractions. A microsplitter was used to obtain a sample of approximately 300 individuals from the >150 μm fraction for a census of the planktic foraminifer assemblage. Foraminifers in the census split were sorted by species and glued onto a standard 60-square micropaleontology slide. Planktic foraminifers were identified following the taxonomy of Parker (1962, 1967), as modified by Poore (1979). Benthic foraminifers were grouped in a single category. Abundances of planktic foraminifers are reported as percent of the total planktic foraminifer assemblage. Chronologies for the cores were based on accelerator mass spectrometry (AMS) 14C dates on mixed planktic foraminifers. Samples of 10–20 mg were converted to graphite targets at the U.S. Geological Survey in Reston, Virginia. AMS 14C ages were determined at the Center for Accelerator Mass Spectrometry, Lawrence Livermore National Laboratory in Livermore, California. The quoted ages are in radiocarbon years (BP) using the Libby halflife of 5568 yr and a marine reservoir correction of 400 yr (Poore et al. 2004) (Table 22.1). Age models derived from the reservoir-corrected radiocarbon dates are shown in Figure 22.2. These independent age models show a nearly uniform accumulation rate of ~20 cm / kyr for the intervals studied in RC 12–10, Gyre 97–6 PC 20, and MD02–2553. The projected ages for the tops of RC 12–10 and Gyre 97–6 PC 20 are ~600 and ~100 kyr, respectively, indicating that modern surface sediments were not recovered in the cores. The projected age for the top of MD02–2553 is 45 yr, but the highest sample available from the core is at a core depth of 5 cm. Sample spacing of the RC 12–10 record varies between 50 to 100 yr but is generally at 50 yr in the upper part of the core (last 7 kyr). Sample spacing in Gyre 97–6 PC 20 is about 50 yr. Average sample spacing in MD02–2553 is 30 yr. The ages reported in this study are in radiocarbon years. Exceptions are noted in the text.

Planktic Foraminiferal Census Data Census data for individual cores are in Appendices 1–3. Foraminifers are abundant and generally moderately well- to well-preserved in samples from core RC 12–10 (Appendix 1). However, the top 44 cm of core RC 12–10 has evidence of increased dissolution in that fragments of planktic foraminifers increase in abundance in counting

370 ~ Poore, Verardo, Caplan, Pavich, and Quinn

Table 22.1. AMS 14C dates (reservoir correction = 400 yr). Depth (cm)

Reservoir corrected age (yr)

Variance (yr)

RC 12-10

1 17 50.5 100.5 135 173 211 255

540 990 2785 4925 6585 8950 11,685 15,310

35 30 60 60 50 40 45 45

Gyre 97-6 PC 20

11.5 40 77.5 100 140

430 1940 3140 4500 6470

30 40 30 40 40

MD02-2553

5.5 29.5 48.5 69.5 110.0 150.5 175.5 200.5 210.0 235.5 250.5 274.5 295.5 310.0 330.5 350.5 375.5 390.5 409.5 424.5

205 960 1440 1970 2865 3620 4150 4675 4755 5465 5920 6505 7390 8120 8220 8770 9505 10,000 10,488 11,055

40 40 40 40 40 40 40 40 40 40 40 40 40 40 40 40 40 40 40 40

Core

splits above 45 cm. Foraminifers are abundant and moderately well- to well-preserved in samples from core Gyre 97–6 PC 20 in the upper 150 cm of the core (Appendix 2). At 150 cm the core has a sharp color change from dark brown to dark gray. Preservation of planktic foraminifers declines below this marker and the abundance of benthic foraminifers increases. The changes seen at 150 cm may represent a slump or dissolution event and therefore the study of Gyre 97–6 PC 20 was limited to the upper 150 cm of the core. Foraminifers are generally well preserved in samples from the upper 350 cm of core MD02– 2553. Below 350 cm depth, preservation of foraminifers in samples declines and fragments are consistently abun-

dant. Counts were stopped at a depth of 367–368 cm due to the declining preservation of the foraminifers (Appendix 3). Reconnaissance examination of samples down to 6 m depth in the core indicates that planktic foraminifers occur sporadically down to about 424 cm and then are absent below 430 cm except for one foraminifer-rich interval at 601–602 cm. Samples checked at 700–702 m, 800–802 cm, 900–902 cm, and 1000–1002 cm were barren of calcareous foraminifers. The stratigraphic distributions of the relative abundance of important species and species groups in the samples are shown in Figures 22.3–22.7. Pink and white varieties of Globigerinoides ruber are consistent components of assemblages from all 3 cores. Globigerinoides ruber (total) accounts for 40–50% of assemblages from RC 12–10 and is 30–40% of assemblages from most samples from MD02–2553 and Gyre 97–6 PC 20. The overall variability of G. ruber appears greater in the last (youngest) 4 ka compared to the older parts of the records. Some taxa had distinctive monotonic stratigraphic trends. For example, the overall relative abundance of Neogloboquadrina dutertrei decreased upsection from the base of the Holocene to the present (Figs. 22.3– 22.5). In contrast, Globorotalia truncatulinoides, Pulleniatina (includes P. obliquloculata and P. finalis), and the G. menardii group (includes G. menardii, G. tumida, and G. ungulata) increased in overall abundance upsection to the present (see Figs. 22.3–22.5). The increase in the G. menardii group is especially evident after about 5 ka. The overall changes in the planktic foraminiferal abundances shown in Figures 22.3–22.5 are similar to previously documented changes in the Gulf of Mexico foraminiferal assemblages related to the transition from glacial to interglacial conditions. Several studies (e.g., Kennett and Huddleston 1972; Flower and Kennett 1995) showed that P. obliquiloculata and the G. menardii group are warm-water (interglacial) indicators that are common in the Holocene of the Gulf of Mexico. The gradual increase of G. truncatulinoides through the Holocene has been documented in a number of cores from the western Gulf of Mexico (Kennett and Huddleston 1972). The data also reveal new features in the abundance variations within the foraminiferal assemblages that add to results of previous studies. Flower and Kennett (1995) found a maximum in the abundance of P. obliquiloculata at 5 ka. These data indicate that the overall abundance of Pulleniatina (counts tabulate P. obliquiloculata and P. finalis as Pulleniatina) increased during the Holocene with no evidence of an abundance peak at 5 ka. An abundance

Planktic Foraminifera and Climate Variability ~ 371

Figure 22.2. Age depth plots for RC 12–10, Gyre 97–6 PC 20, and MD02–2553. Age assignments for samples from MD02–2553 are based on separate models for the upper and lower part of the core. Age assignments for samples from 5.5-cm to 145.5-cm depths are based on the model shown in Figure 22.2C. Age assignments for the samples below 145.5 cm are derived from the model in Figure 22.2D.

maximum in Pulleniatina between 1.6 ka and 1.3 ka in RC 12–10 is not reflected in the Gyre 97 6 PC 20 or MD02– 2553 records. The maximum abundance peak in Pulleniatina in core RC 12–10 occurs in a zone of increased dissolution. Since Pulleniatina is very resistant to dissolution (e.g., Boltovskoy and Toutah 1992), the abundance peak in RC 12–10 could be due, in part, to dissolution. However G. truncatulinoides and the G. menardii group, which are also dissolution resistant, do not have an abundance peak associated with the P. obliquiloculata maximum in RC 12–10. Thus, the Pulleniatina abundance peak in the RC 12–10 record could represent a localized environmental signal in the western Gulf of Mexico. Globorotalia crassaformis disappears from the assemblages in all 3 records at about 7 ka but then briefly reappears in all cores at ~3.2 ka (Fig. 22.6). It is unlikely that the reappearance of G. crassaformis at ~3.2 ka is related to reworking or variable preservation because the reappearance is essentially synchronous and the cores are from different depositional settings and water depths. Previous work (e.g., Kennett and Huddleston 1972) indicated that G. crassaformis is characteristic of glacial intervals in the Gulf of Mexico, and its disappearance in the early Holo-

cene was used to subdivide the Holocene or Z zone of Ericson and Wollin (1968) (Z2 / Z1 boundary of Kennett and Huddleston 1972). Results from this study show that the Z2 / Z1 boundary occurs at about 7 ka and is a reliable datum for dating Holocene sections in the Gulf of Mexico. However, the brief reappearance of G. crassaformis at ~3.2 ka means that the Z2 / Z1 boundary must be used with caution. Globigerinoides sacculifer generally increases in abundance from the base of the Holocene to a broad abundance maximum between 6.5 and 4.5  ka and then decreases in abundance toward the present in all 3 cores (Fig. 22.7). The general increase in G. sacculifer abundance from the base of the Holocene to the mid-Holocene maximum is punctuated by 2 excursions to higher values centered at about 9 and 8.2 ka (cores RC 12–10 and MD02–2553). The millennial-scale pattern of variability is very similar in character and timing between cores (see Fig. 22.3). Other studies (Poore et al. 2003; Poore et al. 2004; Poore et al. 2005) indicate that the G. sacculifer relative-abundance record is an important proxy for climate and environmental changes in the Gulf of Mexico region.

Figure 22.3. Relative abundance variations of selected planktic foraminifers in RC 12–10.

Figure 22.4. Relative abundance variations of selected planktic foraminifers in Gyre 97–6 PC 20.

Figure 22.5. Relative abundance variations of selected planktic foraminifers in MD02–2553.

Figure 22.6. Comparison of the stratigraphic occurrence of Globorotalia crassaformis in cores RC 12–10, Gyre 97–6 PC 20, and MD02–2553. The disappearance and then brief reappearance of G. crassaformis in each core at ~7 ka and ~3 ka is synchronous within the errors of the 14C-derived chronologies. The brief occurrence of G. crassaformis centered at 2 ka in MD02–2553 appears to be separate from the occurrence at ~3 ka and is not seen in the other 2 cores. Bold arrow near 7 ka marks position of the Z2 / Z1 boundary of Kennett and Huddleston (1972). See text for discussion.

374 ~ Poore, Verardo, Caplan, Pavich, and Quinn

Figure 22.7. Comparison of the relative abundance variations of Globigerinoides sacculifer in cores RC 12–10, Gyre 97–6 PC 20, and MD02–2553 and the oxygen-isotope record from ostracods in sediments from Lake Miragoane, Haiti. Each record is controlled by an independent 14C-age model. δ18Ο values from ostracods in Lake Miragoane samples reflect changes in the hydrologic balance, which are manifest as variations in the precipitation / evaporation (P / E) ratio (Hodell et al. 1991). Lower (more negative) δ18Ο values indicate a P / E increase, which in turn indicates higher lake levels (Hodell et al. 1991).

Globigerinoides sacculifer and the Intertropical Convergence Zone Globigerinoides sacculifer is a surface-dwelling species common in the tropical Atlantic and Caribbean. Foraminifer assemblages in core-top samples from the Caribbean usually have greater than 15% G. sacculifer (Kipp 1976). In the Gulf of Mexico, the distribution of G. sacculifer is associated with the Loop Current (Brunner 1979), which brings warm, tropical Caribbean surface waters into the Gulf of Mexico (Fig. 22.1). Globigerinoides sacculifer is very abundant in the core-top assemblage from the southeastern Gulf of Mexico, where the warm Loop Current and Caribbean surface waters are present all year (see Poore et al. 2003, their Fig. 5). The abundance of G. sacculifer in the remaining Gulf of Mexico core-top assemblages declines toward the west and north where tropical SSTs are seasonal. The abundance of G. sacculifer is lowest in areas near the Texas and Louisiana coast where the influence of tropical Caribbean water is minimal. Several studies have shown that changes in incoming

solar insolation related to changes in Earth’s orbit around the sun caused the average position of the ITCZ to migrate during the Holocene (Hodell et al. 1991; Haug et al. 2001). In the early Holocene, increased summer insolation resulted in warming of the Northern Hemisphere and northward movement of the average position of the ITCZ. As the average position of the ITCZ moved northward, southeasterly winds were enhanced resulting in increased precipitation along the northern coast of Venezuela, as monitored using runoff records in the Cariaco basin (Haug et al. 2001) and lake-level records at Lake Miragoane in Haiti (Hodell et al. 1991). After about 5 ka, decreasing summer insolation in the Northern Hemisphere and resulting southward movement of the average position of the ITCZ caused a decline in southeasterly winds. Thus, precipitation in Haiti and along the north coast of Venezuela declined (Hodell et al. 1991; Haug et al. 2001). Poore et al. (2004) and Poore et al. (2005) reasoned that changes in surface-water circulation in the Gulf of Mexico parallel the changes in atmospheric circulation inferred from the Lake Miragoane and Cariaco basin records and con-

Planktic Foraminifera and Climate Variability ~ 375

cluded that G. sacculifer in assemblages from the western and northern Gulf of Mexico reflect changes in the average position of the ITCZ. Northward migration of the ITCZ and the resulting enhancement of the summer circulation pattern are indicated by increased abundance of G. sacculifer in faunal assemblages at the beginning of the Holocene leading up to maximum abundances between 6.5 and 4.5 ka (Fig. 22.7). After reaching maximum values in the mid-Holocene, the abundance of G. sacculifer declines toward the present as the average position of the ITCZ migrated back toward the equator. The excursions to high values of G. sacculifer near 9.0 and 8.2 ka coincide almost exactly with brief intervals of elevated lake levels in Lake Miragoane, Haiti (Hodell et al. 1991) (Fig. 22.7). The 9.0- and 8.2-ka excursions present in the Gulf of Mexico and Lake Miragoane records presumably represent brief but regionally significant northward excursions of the ITCZ during the early Holocene. Future work is needed to confirm the presence of the 9.0and 8.2-ka events in additional proxy records from the Caribbean and Gulf of Mexico area.

Century-Scale Variability Inspection of the G. sacculifer abundance plots in Figure 22.7 reveals century-scale cycles superimposed on the long-term Holocene trends. The cycles are most evident in the more highly resolved MD02–2553 record and are also clearly developed in the RC 12–10 and Gyre 97–6 PC 20 records between ~6 ka and 2 ka. Analyses of a number of late Pleistocene and Holocene climate-proxy time series, including the Gulf of Mexico G. sacculifer abundance records, reveal statistically significant century-scale cycles that are very similar to concentrations of variance in the 14C-production record. For example, a 208-yr cycle is prominent in the 14C-production record, and cycles near

200 yr are also present in highly resolved climate-proxy records from a variety of regions (Table 22.2). Since variability in 14C production is related to changes in solar output, the occurrence of similar variability in climate-proxy and 14C-production records has been used as evidence linking solar variability to century-scale climate change (e.g., see discussions in Bond et al. 2001). Poore et  al. (2004) concluded that century-scale variability in G. sacculifer abundance was related to solar-forced changes in the average position of the ITCZ. Warmer intervals caused by increased solar output would result in a more northerly average position of the ITCZ and higher G. sacculifer abundances in Gulf of Mexico foraminifer assemblages. Cooler intervals caused by diminished solar output would result in a more southerly average position of the ITCZ and lower G. sacculifer abundances. Variability in the average position of the ITCZ should influence monsoon circulation in the southwestern United States. Northward migration of the average position of the ITCZ would increase monsoon precipitation, and migration of the ITCZ toward the equator would decrease monsoon precipitation. In Figure 22.8 the G. sacculifer record from MD02–2553 is compared with the long tree-ring record from El Malpais National Monument on the southern periphery of the Colorado Plateau in west-central New Mexico (see Fig. 22.1). The El Malpais record is important because it is continuous from 136 BCE (2136 yr BP) to 1992, and variations in tree-ring width in the record are highly correlated with annual (water-year) precipitation (Grissino-Mayer 1996). Because precipitation in New Mexico occurs predominantly in midsummer to fall (June to October, Higgins et al. 1997, their Fig. 4), the El Malpais record is considered to be particularly sensitive to variations in the southwest monsoon. The MD02–2553 and El Malpais records shown in Figure 22.8 are plotted on their own indepen-

Table 22.2. Climate records showing significant variability near 200-yr cycle. Proxy

Location

ice-rafted grains sediment density G. bulloides % G. sacculifer % G. ruber δ18O speleothem δ18O sediment grayscale sediment clay %

sub-polar North Atlantic Yucatan Peninsula Cariaco basin Gulf of Mexico Arabian Sea Dongge Cave, southern China Santa Barbara basin Palmer Deep Antarctic

Interval (ka)

Period

Resolution

0–12 0–2.6 1–10 2.8–7.4 6–10 0–9 3.8–9.9 0–8.7

225 208 200 212 220 206 192 200

30–50 yr 11–14 yr ~33 yr 50 yr ~20 yr 5 yr seasonal 3–5 yr

Source Bond et al. 2001 Hodell et al. 2001 Peterson et al. 1991 Poore et al. 2003 Staubwasser et al. 2002 Wang et al. 2005 Schaaf and Thurow 1998 Warner and Domack 2002

376 ~ Poore, Verardo, Caplan, Pavich, and Quinn

dent age model. The tree-ring record is in calendar years so the OxCal program was used to calibrate the AMS 14 C-controlled chronology for MD02–2553 to calendar years (= calibrated years). Low-frequency variations were removed from the MD02–2553 record to match the treering record, which had low-frequency trends removed during standard processing. Previous studies showed that the general southward shift of the average position of the ITCZ and associated decrease in monsoon rainfall in the Caribbean region over the past 5 ka coincided with an overall decrease of G. sacculifer in Gulf of Mexico sediments (Hodell et al. 1991; Haug et al. 2001; Poore et al. 2004). Thus Poore et al. (2005) inferred that the multicentury drought centered at about 1600 yr BP in the El Malpais record coincided with the multicentury interval of negative residuals (low G. sacculifer abundance) in the de-trended MD02–2553 record. The 2 features were matched in Figure 22.8 by uniformly shifting the MD02– 2553 record 60 yr, which is well within the errors involved in 14C dating and conversion of 14C to calendar years. Inspection of Figure 22.8 reveals substantial overall similarity between the 2 records. For example, precipita-

tion maxima near 1400 yr BP and 900 yr BP are prominent in both records. Correlation of the records between 1700 yr BP and 1400 yr BP is high (r 2 = 0.66). Major dry excursions seen in the El Malpais record at ~2100 yr BP, ~1950 yr BP, ~1000 yr BP, and ~400 yr BP correspond closely with minima in the MD02–2553 record. Over some intervals, for example between 900 yr BP and 500 yr BP, the MD02–2553 record appears slightly distorted with respect to the El Malpais record, as if the marine record had undergone minor stretching and compression. The apparent distortion could easily be explained by minor mechanical disturbance during the coring process or by small changes in sediment-accumulation rates in the marine core. Alternatively, other variables in the climate system, such as a strong El Niño–Southern Oscillation event, could have modified seasonal distribution of precipitation in New Mexico for a time. Although additional highly resolved records from the Gulf of Mexico are needed to verify the details of the MD02–2553 record, the overall similarity between the MD02–2553 and the El Malpais records confirms the interpretation that variations in G. sacculifer abundance in sediments of the north-

Figure 22.8. Comparison of smoothed time series of standard tree-ring width index record from El Malpais National Monument in west-central New Mexico with residual from de-trended relative abundance record of Globigerinoides sacculifer in core MD02–2553 from Pigmy Basin in the northern Gulf of Mexico (see Fig. 22.1). MD02–2553 record has been de-trended to remove low-frequency variability. Timescale is in calendar years before 2000 (see Poore et al. 2005). MD02–2553 record has been uniformly shifted 60 yr younger with respect to the El Malpais chronology (see text for discussion). Larger values of tree-ring width index represent increased annual precipitation. Larger values of G. sacculifer residual values represent more northerly average position of the ITCZ.

Planktic Foraminifera and Climate Variability ~ 377

ern Gulf of Mexico are linked to century-scale changes in the average position of the ITCZ and provide a reliable proxy for century-scale variations in the southwest monsoon.

MD02–2553. We thank Ellen Roosen for assistance sampling Gyre 97–6 PC 20 and Rusti Lotti for assistance in sampling RC 12–10. Lisa Osterman, Barbara Lidz, and Peter Swarzenski provided constructive comments that improved the manuscript.

Summary and Conclusions Census of planktic foraminiferal assemblages from 3 AMS 14 C-dated cores provides basic data for determining Holocene climate and environmental variability in the Gulf of Mexico region. Distinct trends and millennial-scale variations in the census data from the 3 separate, independently dated cores are very similar in timing and magnitude. The coherent pattern of the census data among cores suggests the records represent a reliable indicator of the planktic foraminifer populations in the Gulf of Mexico. Any distortions of the records due to dissolution, reworking, or missing sections are minor. Results from this study show that the regional extinction of Globorotalia crassaformis in the Gulf of Mexico, and thus the Z2 / Z1 boundary of Kennett and Huddleston (1972), occurs at 7 ka. However, from the census data, G. crassaformis briefly reappears in the Gulf of Mexico for an interval of several hundred years at about 3.2 ka. Thus, identification of Z2 and Z1 must be done with caution. The relative abundance of the G. sacculifer in Gulf of Mexico foraminiferal assemblages appears to be a reliable proxy for changes in the average position of the ITCZ. The census data confirm previous work indicating that the average position of the ITCZ reached its northern extent in the mid-Holocene and has been migrating southward from the mid-Holocene to the present. Two brief northward excursions of the ITCZ at about 9 and 8.2 ka are evident in the lake-level record from Lake Miragoane, Haiti, and in Gulf of Mexico cores RC 12–10 and MD02–2553. Higher-frequency (century-scale) changes in the average position of the ITCZ and associated changes in the southwest monsoon are also reflected in variation in the G. sacculifer relative-abundance data and appear to be related to changes in solar output.

Acknowledgments We thank Sarah Medley and Amy Spaziani for technical assistance during the study. We thank Dave Hodell for data from Lake Miragoane, Dave Twichell for providing access to Gyre 97–6 PC 20, and Ben Flower for access to

Taxonomic Notes The following list is comprised of species and groupings of species mentioned in the text and in Appendices 1–3; notes are provided in some instances for clarification. In general, species concepts follow those of Parker (1962, 1967) and Poore (1979). Globigerina bulloides (d’Orbigny). Aside from the early Holocene in RC 12–10, G. bulloides is usually a very minor component of the assemblages. Globigerina calida (Parker). It is sometimes difficult to separate tightly coiled specimens of Globigerina calida from specimens of Globigerinella aequilateralis that do not show clear planispiral coiling. By convention, transitional individuals are assigned to G. calida if the extra-umbilical extent of the aperture does not reach the midpoint between the umbilical and spiral side of the test when seen in side view. For example, the individual illustrated in Figure 24 on Plate 2 of Parker (1962) would be assigned to G. calida, not to Globigerinella aequilateralis. Globigerina digitata (Brady) Globigerina falconensis (Blow) Globigerinella aequilateralis (Brady). See comments for Globigerina calida. Globigerinita glutinata (Egger). Globigerinita naparimaensis Bronnimann and Globigerina juvenilis Bolli are included in Globigerinita glutinata. Globigerinoides conglobatus (Brady) Globigerinoides ruber (d’Orbigny). Globigerinoides ruber occurs in pink and white varieties. The development of pink or red coloration may be related to photosynthetic symbionts living in the protoplasm of the foraminifer. By convention, a specimen is tabulated as pink even if only one chamber is pink or red. In general, the white form of G. ruber is more abundant than the red variety, although exceptions do occur (e.g., in RC 12–10 samples 176–178 cm and 180–182 cm; see Appendix 1). Together, the red and white varieties of G.

378 ~ Poore, Verardo, Caplan, Pavich, and Quinn

ruber are consistently the most abundant component in the samples (Appendices 1–3). Globigerinoides sacculifer (Brady). Forms with and without the final sac-like final chamber are included in Globigerinoides sacculifer. Thus, forms placed in Globigerinoides quadrilobatus (d’Orbigny) and G. trilobus (Reuss) are tabulated as G. sacculifer. Globorotalia crassaformis (Galloway and Wissler) Globorotalia inflata (d’Orbigny) Globorotalia menardii (Parker, Jones, and Brady). Census data show that Globorotalia menardii, G. tumida, and G. ungulata (the G. menardii group) are rare and sporadic in foraminiferal assemblages from the earliest part of the Holocene. The G. menardii group becomes a common and persistent component of the foraminiferal assemblages in the youngest part of the Holocene (see Figs. 22.3–22.5). Globorotalia truncatulinoides (d’Orbigny) Globorotalia tumida (Brady) Globorotalia ungulata Bermudez. Large, encrusted representatives of Globorotalia ungulata can be difficult to separate consistently from G. tumida. Neogloboquadrina dutertrei (d’Orbigny) Orbulina universa (d’Orbigny) Pulleniatina spp. Includes P. obliquiloculata (Parker and Jones) and P. finalis (Banner and Blow). Most of the individuals are referable to P. obliquiloculata, but a few individuals referable to P. finalis are present in many samples. Other. This category includes unidentified individuals as well as taxa that occur sporadically in very low abundances. Candeina nitida (d’Orbigny), Globigerina rubescens (Hofker), Globorotalia scitula (Brady), and Neogloboquadrina dutertrei–N. pachyderma (Ehrenberg) intergrades (N. dupac of Poore 1979) are taxa that are rare and sporadic in the samples.

References Boltovskoy, E., and V. I. Toutah. 1992. Preservation index and preservation potential of some foraminiferal species. Journal of Foraminiferal Research 22:267–73. Bond, G., B. Kromer, J. Beer, R. Muscheler, M. N. Evans, W. Showers, S. Hoffmann, R. Lotti-Bond, I. Hajdas, and G. Bonani. 2001. Persistent solar influence on North

Atlantic climate during the Holocene. Science 294: 2130–36. Brown, P., J. P. Kennett, and B. L. Ingram. 1999. Marine evidence for episodic Holocene megafloods in North America and the northern Gulf of Mexico. Paleoceanography 14:498–510. Brunner, C. A. 1979. Distribution of planktonic foraminifera in surface sediments of the Gulf of Mexico. Micropaleontology 25:325–35. ———. 1982. Paleoceanography of surface waters in the Gulf of Mexico during the Late Quaternary. Quaternary Research 17:105–19. Elderfield, H., and G. Ganssen. 2000. Past temperature and δ18O of surface ocean waters inferred from foraminiferal Mg / Ca ratios. Nature 405:442–45. Ericson, D. B., and G. Wollin. 1968. Pleistocene climates and chronology in deep-sea sediments. Science 162:1227–34. Flower, B. P., and J. P. Kennett. 1990. The Younger Dryas cool episode in the Gulf of Mexico. Paleoceanography 5:949–61. ———. 1995. Biotic responses to temperature and salinity changes during the last deglaciation, Gulf of Mexico. Pp. 209–20 in Effects of Past Global Changes on Life. Washington, D.C.: National Academy of Sciences. Grissino-Mayer, H. D. 1996. A 2129-year reconstruction of precipitation for northwestern New Mexico USA. Pp. 191–204 in J. S. Deat, D. M. Meko, and T. W. Swetnam (eds.), Tree Rings, Environment, and Humanity. Tucson, Ariz.: University of Arizona. Haug, G. H., K. A. Hughen, D. M. Sigman, L. C. Peterson, and U. Röhl. 2001. Southward migration of the Intertropical Convergence Zone through the Holocene. Science 293:1304–308. Higgins, R. W., Y. Yao, and X. L. Wang. 1997. Influence of the North American monsoon system on the U.S. summer precipitation regime. Journal of Climate 10:2600–622. Hodell, D. A., J. H. Curtis, G. A. Jones, A. Higuera-Gundy, M. Brenner, M. W. Binford, and K. T. Dorsey. 1991. Reconstruction of Caribbean climate change over the past 10,500 years. Nature 352:790–93. Hodell, D. A., B. Mark, J. H. Curtis, and T. Guilderson. 2001. Solar forcing of drought frequency in the Maya Lowlands. Science 292:1–4. Imbrie, J., and N. G. Kipp. 1971. A new micropaleontological method for quantitative paleoclimatology: application to a Late Pleistocene Caribbean core. Pp. 71–181 in K. K. Turekian (ed.), Late Cenozoic Glacial Ages. New Haven, Conn.: Yale University Press. Kennett, J. P., K. Elmstrom, and N. Penrose. 1985. The last deglaciation in Orca Basin, Gulf of Mexico: high-resolution

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planktonic foraminiferal changes. Palaeogeography, Palaeoclimatology, Palaeoecology 50:189–216. Kennett, J. P., and P. Huddleston. 1972. Late Pleistocene paleoclimatology, foraminiferal biostratigraphy and tephrochronology, western Gulf of Mexico. Quaternary Research 2:38–69. Kipp, N. G. 1976. New transfer function for estimating past sea-surface conditions from sea-bed distribution of planktonic foraminiferal assemblages in the North Atlantic. Pp. 3–41 in R. M. Cline and J. D. Hayes (eds.), Investigations of Late Quaternary Paleoceanography and Paleoclimatology. Boulder, Colo.: Geological Society of America. Malmgren, B., and J. P. Kennett. 1976. Principal component analysis of Quaternary planktic foraminifera in the Gulf of Mexico: paleoclimatic applications. Marine Micropaleontology 1:299–306. Mayle, F. E., R. Burbridge, and T. J. Killeen. 2000. Millennialscale dynamics of southern Amazonian rain forests. Science 290:2291–94. Metcalfe, S. E., S. L. O’Hara, M. Caballero, and S. J. Davies. 2000. Records of Late Pleistocene-Holocene climatic change in Mexico-a review. Quaternary Science Reviews 19:699– 721. Müller-Karger, F. E., J. J. Walsh, R. H. Evans, and M. B. Meyers. 1991. On the seasonal phytoplankton concentration and sea surface temperature cycles of the Gulf of Mexico as determined by satellites. Journal of Geophysical Research 96:12,645–65. Oglesby, R. J., K. A. Maasch, and B. Saltzman. 1989. Glacial melt water cooling of the Gulf of Mexico: GCM implications for Holocene and present-day climates. Climate Dynamics 3:115–33. Parker, F. L. 1962. Planktonic foraminiferal species in Pacific sediments. Micropaleontology 8:219–54. ———. 1967. Late Tertiary biostratigraphy (planktonic foraminifera) of tropical Indo-Pacific deep-sea cores. Bulletins of American Paleontology 52:115–208. Peixóto, J. P., and A. H. Oort. 1983. The atmospheric branch of the hydrological cycle and climate. Pp. 5–65 in A. Street-Perrott, M. Beran, and R. Ratcliffe (eds.), Variations in the Global Water Budget. Boston, Mass.: D. Reidel Publishing. Petersen, K. L. 1994. A warm and wet little climatic optimum and a cold and dry Little Ice Age in the southern Rocky Mountains, U.S.A. Climatic Change 26:243–69.

Peterson, L. C., J. T. Overpeck, N. G. Kipp, and J. Imbrie. 1991. A high-resolution Late Quaternary upwelling record from the anoxic Cariaco Basin, Venezuela. Paleoceanography 6:99–119. Poore, R. Z. 1979. Oligocene through Quaternary planktic foraminiferal biostratigraphy of the North Atlantic: DSDP Leg 49. Pp. 447–517 in B. P. Luyendyk, J. R. Cann, et al. (eds.), Initial Reports of the Deep Sea Drilling Project, Vol. 49. Washington, D.C.: U.S. Government Printing Office. Poore, R. Z., H. J. Dowsett, and S. Verardo. 2003. Millennial- to century-scale variability in Gulf of Mexico Holocene climate records. Paleoceanography 18:26-1–26-13. Poore, R. Z., T. M. Quinn, and S. Verardo. 2004. Century-scale movement of the Atlantic Intertropical Convergence Zone linked to solar variability. Geophysical Research Letters 31:L12,214. Poore, R. Z., M. J. Pavich, and H. D. Grissino-Mayer. 2005. Record of the North American southwest monsoon from Gulf of Mexico sediment records. Geology 33:209–12. Ravelo, A. C., R. G. Fairbanks, and S. G. H. Philander. 1990. Reconstructing tropical Atlantic hydrography using planktonic foraminifera and an ocean model. Paleoceanography 5:409–31. Schaaf, M., and J. Thurow. 1998. Two 30,000-year highresolution greyvalue time series from the Santa Barbara Basin and the Guaymas Basin. Pp. 101–10 in A. Cramp, C. J. MacLeod, S. V. Lee, and E. J. W. Jones (eds.), Geological Evolution of Ocean Basins: Results from the Ocean Drilling Program. Special Publication 131. Bath, U.K.: The Geological Society. Tang, M., and E. R. Reiter. 1984. Plateau monsoons of the Northern Hemisphere: a comparison between North America and Tibet. Monthly Weather Review 112:617–37. Staubwasser M., F. Sirocko, P. Grootes, and H. Erlenkeuser. 2002. South Asian monsoon climate change and radiocarbon in the Arabian Sea during early and middle Holocene. Paleoceanography 17:15.1–15.12. Warner, N. R., and E. W. Domack. 2002. Millennial- to decadal-scale paleoenvironmental change during the Holocene in the Palmer Deep, Antarctica, as recorded in particle size analysis. Paleoceanography 7:5.1–5.14. Wang, Y., H. Cheng, R. L. Edwards, Y. He, X. Kong, Z. An, J. Wu, M. J. Kelly, C. A. Dykoski, and X. Li. 2005. The Holocene Asian monsoon: links to solar changes and North Atlantic climate. Science 308:854–57.

␥23

Over 300 Years of Anthropogenic and Naturally Induced Low- Oxygen Bottom-Water Events on the Louisiana Continental Shelf Lisa E. Osterman, Richard Z. Poore, Peter W. Swarzenski, David Hollander, and R. Eugene Turner

Oxygen-depleted subsurface waters occur on the Louisiana continental shelf when the uptake of oxygen exceeds its resupply. When oxygen concentrations fall below 2 mg / L, it is operationally defined as hypoxic. Measurements of Louisiana continental shelf waters demonstrate that the size of the hypoxic zone has increased since 1985 (Rabalais et al. 1999; CENR 2000; Rabalais and Turner 2001) (Fig. 23.1). The likelihood of hypoxia occurring is increased with enhanced nutrient loading and the onset of greater water-column stratification (van der Zwaan 2000; Rabalais 2002). Increased nutrients can stimulate marine surface phytoplankton blooms whose organic matter, upon death, may sink to the bottom and decay. The modern recurrent hypoxia on the Louisiana shelf and expansion of the area of the hypoxic zone has been attributed to the increased use of fertilizer in the Mississippi River basin (Dinnel and Bratkovich 1993; Nelson et al. 1994; Rabalais et al. 1994; Rabalais et al. 1996; Rabalais et al. 1999; Goolsby et al. 2001). The increased frequency and expansion of hypoxic conditions has become an important economic and environmental issue to commercial and recreational fisheries in this region of the Gulf of Mexico (Malakoff 1998; Diaz and Solow 1999). Beginning in 1985, systematic summertime surveys have measured the oxygen content of subsurface waters across the Louisiana shelf (Rabalais et al. 1999). Because the hypoxia exists on an open shelf as opposed to a closed basin, the temporal and spatial extent fluctuates throughout the year but is concentrated in the warmer summer months when thermal stratification inhibits water column

overturning. Therefore, the zone is expressed as the past frequency of hypoxia occurrence at a particular location since the initiation of shelfwide monitoring (Rabalais et al. 1999) (Fig. 23.1). The survey data demonstrate that the size of the area affected by hypoxia has been increasing since 1985. However, no routine shelfwide measurements of the extent of subsurface-water oxygen content were made before 1978 (Rabalais et al. 1999; Turner et al. 2005). In this chapter, we discuss the use of low-oxygen– tolerant benthic foraminifers and geochemical measurements in sediment cores from the Louisiana shelf as proxies for extending the observed record of low-oxygen bottom-water conditions back through time and space.

Low- Oxygen Faunal Proxy Both living and fossil calcareous benthic foraminifers have proven to be valuable tools for assessing oxygen conditions of bottom water (Bernhard 1986; van der Zwaan and Jorissen 1991; Kaiho 1994) and as a proxy for bottom-water hypoxia in the Santa Barbara Basin, Adriatic Sea, and elsewhere (Bernhard et al. 1997; Bernhard and Sen Gupta 1999; Duijnstee et al. 2004). Only a few studies have investigated the faunal sedimentological record near or within the Louisiana hypoxia zone. Blackwelder et al. (1996) investigated core BL10 from near the Mississippi delta (Fig. 23.1). Core BL10 (sediment accumulation rate = 0.55 cm / yr), collected in 1993, has a 210Pb chronology covering ~90 years and provides evidence for 381

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Figure 23.1. Location of cores discussed in this chapter. Black circles indicate USGS cores reported in this chapter. Stars indicate the location of 2 previously reported cores, BL10 (Blackwelder et al. 1996) and G27 (Sen Gupta et al. 1996; Platon et al. 2005). Location of hypoxia zone frequency taken from CENR (2000), based on summertime measurements done by Rabalais since 1985.

an increasing trend in hypoxia in which an older assemblage of agglutinated, miliolid, and hyaline foraminifers is replaced upsection by a foraminiferal assemblage of lowoxygen–tolerant hyaline species, including Epistominella vitrea, Buliminella morgani, and Pseudononion atlanticum (Blackwelder et al. 1996). This faunal change is accompanied by a decrease in diversity (H′) from 1.0 at 1900 to 0.7 near the top of the core (Nelson et al. 1994). These diversity and species changes are interpreted as an indication of increasing hypoxia during the last 90 years caused by increased fertilizer use and transport to the Gulf of Mexico via Mississippi River outflow. Sen Gupta et al. (1996), Platon and Sen Gupta (2001), and Platon et al. (2005) discussed the Ammonia–Elphidium (A–E) hypoxia index from 11 box core records collected in the easternmost region of the Louisiana hypoxia zone. The A–E index = number of Ammonia parkensoniana / (number of Ammonia parkensoniana + number of Elphidium excavatum). Sediment accumulation rates based on 210Pb vary between 0.1 and 2.0  cm / yr. Most cores had evidence of increasing hypoxia during the last 40 years as monitored by the A–E hypoxia index, but only 3 cores contained records greater than 70 years. In part, this is because the A–E hypoxia index is based on taxa that inhabit marginal marine and inner-shelf environments less than ~20 m water depth where longer sediment records are more difficult to obtain. The A–E hypoxia index is not considered reliable in greater than ~30 m mean water depth (Platon and Sen Gupta 2001). More recently, Osterman (2003) analyzed 74 continental shelf core-top samples from both the Texas and

Louisiana continental shelf and slope and showed that the distribution of 3 benthic foraminiferal species can be used to characterize the hypoxia zone. The cumulative percentage of 3 foraminifers, named the PEB index (P. atlanticum, E. vitrea, and B. morgani), was higher in the surface-sediment samples collected in the Louisiana hypoxia zone than elsewhere along the Texas–Louisiana shelf (Osterman 2003). Benthic foraminiferal assemblages in surface samples from the eastern part of the hypoxia zone (Platon and Sen Gupta 2001; Platon et al. 2005) confirm that the PEB index species are the most common living foraminiferal species in the Louisiana hypoxia zone. The PEB species also live in slightly deeper water (30– 70 m water depth) than the taxa used in the A–E hypoxia index (Sen Gupta et al. 1996), and thus the PEB index can extend the application of foraminiferal hypoxia proxies over a wider area of the continental shelf. The PEB species are opportunists that inhabit nutrient-rich environments, tolerate low-oxygen conditions, and are believed to be epifaunal (Blackwelder et al. 1996; Austin and Evans 2000; Gooday and Hughes 2002) or migrate to epifaunal habitats in nutrient-rich, low-oxygen environments (Jorissen et al. 1992; Ernst and van der Zwaan 2004). Total loss of oxygen (anoxia) will result in death for many foraminiferal species, but a few opportunistic, lowoxygen–tolerant species have adapted to withstand short intervals (days to weeks) of anoxia (Bernhard and Alve 1996; Bernhard et al. 1997; Moodley et al. 1998; Bernhard and Sen Gupta 1999; Duijnstee et al. 2004). Several mechanisms have been suggested to explain the tolerance of some benthic foraminifers to low-oxygen conditions including symbiont farming or “hibernation.” Symbiont farming (the sequestration of chloroplasts) has been observed in crenulated apertures of facultative anaerobes living in the anoxic environment of the Santa Barbara Basin (Bernhard and Bowser 1999). Platon and Sen Gupta (2001) suggested that the PEB species may also be facultative anaerobes, harboring symbionts capable of producing oxygen, food, or assisting in biochemical processes, which would enable these foraminifers to withstand hypoxic episodes. The exact mechanism used by the PEB species to survive low-oxygen conditions is unknown. Given the highly variable temporal and spatial distribution of seasonal hypoxia on the open Louisiana shelf (Rabalais et al. 1999), we conclude that hypoxia-tolerant species will continue to live through hypoxic episodes and reproduce, whereas other less-tolerant species do not, resulting in a relative increase of the PEB species during recurrent (seasonal or

Low-Oxygen Bottom-Water Events ~ 383

annual) low-oxygen episodes. Our working model is that the increased relative abundance of PEB species in sediment cores reflects times of development of low-oxygen bottom-water events and even hypoxia on the Louisiana continental shelf. Remember that in older records, more removed from the modern analog, additional sedimentological and biological conditions must be considered to assess the environmental conditions.

work for the 2 MRJ03 cores and PE0305-GC1 (=GC1) can be found in Swarzenski et al. (2006). Precision in the activities of 210Pb, 239,240Pu, and 226Ra was typically better than 5%. A separate set of samples was taken from core GC1 for analyses of carbon content. Concentrations of total carbon and total inorganic carbon (IC) were determined using a UIC carbon coulometer. A known quantity of each sediment sample (approximately 40 mg) was weighed and placed in a porcelain boat into the carbon coulometer oven, which combusted the sample at 970 °C in the presence of excess oxygen. All of the carbon was oxidized into carbon dioxide to determine the total carbon concentration. A known quantity of each sediment sample was weighed, placed into a glass sample flask, and then acidified in a heated reaction vessel using 5 mL of 2 mol perchloric acid to release all the inorganic carbon. Total organic carbon (TOC) was calculated as the difference between the total carbon (released by combustion) minus total inorganic carbon (generated by acidification). A single AMS 14C date of 4195 ± 35 years on mixed benthic foraminifers was obtained in the lower part of gravity core GC1 (152–157 cm). The graphite target for the date was processed at the 14C laboratory of the U.S. Geological Survey in Reston, Virginia, and the 14C age was determined at the Center for Accelerator Mass Spectrometry (CAMS), Lawrence Livermore National Laboratory, Livermore, California. The quoted age is in radiocarbon years (BP) using the Libby half-life of 5568 years, and no correction has been applied (Fig. 23.4).

Materials and Methods Our results are based on study of a series of box and gravity cores recovered from the Louisiana shelf to the Mississippi Delta (Figs. 23.1, 23.2). In general, cores were sampled every centimeter for analyses of benthic foraminiferal assemblages and geochronological tracers. Information about core collection, sample processing, benthic foraminifer counts, taxonomic notes, figure references, and the PEB index are reported elsewhere (Osterman et al. 2004; Osterman et al. 2005; Swarzenski et al. 2006). Development of the PEB index is discussed in Osterman (2003) and Osterman et al. (2005). Geochronology for the top portions of our cores was estimated using multiple radioactive tracers including excess 210Pb, 137Cs, and 239, 240Pu isotopes (Baskaran et al. 1996) (Fig. 23.3). Core PE0305-BC1 sediments were analyzed using an integrated gamma-spectroscopy system (Nittrouer et al. 1979). A detailed description of the methods utilized in developing a geochronological framePE0305-BC1(47 mwd) a) 2000

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Figure 23.2. Increasing trend of low-oxygen bottom water over the last 50 years seen in 6 Louisiana shelf cores with excess 210Pb chronologies. (A–D) PEB index values for 4 cores reported in this chapter (error bars on the PEB ratio average 3%). (E) PEB values for BL10 calculated from foraminiferal data in Blackwelder et al. (1996, Table 4). F) A-E Index values for G27 are modified from Sen Gupta et al. (1996, Fig. 4). Note that 2 cores (GC1 and G27) contain longer records, which are not shown in this figure. Letters C, D, and E are PEB peaks that appear in more than one core and are also noted in Fig. 23.4.

384 ~ Osterman, Poore, Swarzenski, Hollander, and Turner

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0.5

210

Results and Discussion The PEB index in 4 sediment cores with 210Pb-controlled dating is shown in Figure 23.2. Chronologies plotted on the figure are derived from the 210Pb data (Fig. 23.3) and assume constant accumulation rates.

PEB—The Last 150 Years The cores had a similar stratigraphic pattern of generally increasing PEB values from 1950 to the present (Fig. 23.2) in all 4 cores. The cores were collected from ~30 to 70 m water depth and collected within an area of about 2500 km2 on the Louisiana shelf. Thus, the PEB data indicate that the post-1950 increase in low-oxygen bottom-water conditions is a robust signal over this part of the Louisiana shelf. The trend towards higher PEB values over the last 50 years is consistent with previous studies that have linked the increased occurrence of hypoxia on the Louisiana shelf to higher nutrient loading associated

1

Pb dpm/g xs

1.5

Figure 23.3. Excess 210Pb (dpm / g) record and age models for the 4 cores reported in this chapter.

with the use of commercial fertilizer in the Mississippi River basin (Nelson et al. 1994). Several brief excursions to higher values of the PEB index in about 1880, 1900, and 1920 suggest low-oxygen bottom-water events or intervals occurred in the late 1800s and early 1900s (labeled C-E in Figs. 23.2, 23.4) (Osterman et al. 2005). The maximum values of the pre-1950 excursions in the PEB are less than the maximum post-1950 PEB values observed near the top of each core, but several of the excursions are present in more than one core, which suggests they record real events. The development of continental shelf hypoxia depends on a variety of factors (timing, duration, and onset of stratification; storm-induced physical mixing of surface waters; or core location and water depth) that could influence the geographic extent, duration, and intensity of hypoxic events. This helps explain why lowoxygen events may not be recorded in all cores. Two additional cores provide supporting evidence for low-oxygen events on the Louisiana shelf. The A–E hypoxia index values in core G-27 (Sen Gupta et al. 1996;

Low-Oxygen Bottom-Water Events ~ 385

PE0305-GC1 a)

0

+1SD 2000

+1SD 15-100 cm

0

Core Description c)

b)

0

A

20

B

40

D E

60

F G

20

20

40

40

60

60

80

80

1700 ? 100

100

1900

80

Estimated Age

Depth (cm)

1800 ?

H I

100

J 120

120

120

K L

140

4195 ± 35 BP

160 0

suspected discontinuity

5

140

140

160

160

10 15 20 25 30 35

PEB Index

0

100 200 300 400 500 600

N/gm

Mollusk Burrows Silt layer Clay

Figure 23.4. Summary of PE0305-GC1. (A) PEB index values for core. Peaks A, I, J, and L are greater than one standard deviation (SD) of the mean for the entire core, excluding the interval of suspected disconformity (120–137 cm). Peaks B, D, E, F, G, and H are >1 SD in the core interval from 15–100 cm. Peaks B, D, and E are also seen in Figure 23.2. Peaks D-G are discussed in Osterman et al. (2005). Location of 14C date (152–157 cm) is also shown. (B) The number of all benthic foraminifers per gram of dry sediment (N / gm). See Osterman et al. (2006) for more information. (C) Core description from X-radiographs.

Platon et  al. 2005) show an increasing trend of lowoxygen bottom water since 1950 and higher A–E values (lower oxygen) around 1880 and 1920 (Figs. 23.1, 23.2). In addition, calculation of the PEB index from the BL10 foraminiferal data (Blackwelder et  al. 1996) indicates increasing PEB values after 1950, as well as a significant (+1 SD of pre-1950 values) increase in the PEB species in approximately 1910 and 1900 (Fig. 23.2). These similar results from 3 different groups of workers, using 2 different methods, indicate that the post-1950 increase in lowoxygen bottom water is a robust signal over a wide area of the Louisiana shelf (Blackwelder et al. 1996; Sen Gupta et al. 1996; and this chapter). Fluctuations in the PEB index before 1950 have been attributed to natural variations in Mississippi River discharge that enhances transport of nutrients and freshwater to nearshore waters (Poore et al. 2001; Osterman et al. 2005). A combination of natural processes as well as possible anthropogenic activities, such as land clearing in the Mississippi basin, may have resulted in the devel-

opment of widespread algal blooms and resulting lowoxygen bottom water on the Louisiana continental shelf before 1950 (Justić et al. 2003; Turner and Rabalais 2003). Our data indicate that the low-oxygen events of the last few decades were more extreme than any that occurred in the previous ~150 years and support the interpretation that the increased use of commercial fertilizer has amplified an otherwise naturally occurring process.

PEB–Before 1850 The PEB record below 50 cm (~1850) in GC1 includes at least 3 minor excursions to elevated PEB values (F-H) and 4 even older excursions to very large PEB values (I-L) that are comparable to values found in the youngest part of the record when the Louisiana hypoxia zone was intensified by the use of commercial fertilizer (Fig. 23.4). However, as we diverge from the modern analog, more care must be taken with the dating and interpretation of these older PEB excursions.

386 ~ Osterman, Poore, Swarzenski, Hollander, and Turner

Age Model The geochronology for the long record in core GC1 is constrained by 210Pbxs-derived age dates in the upper part of the core and one AMS 14C date of ~4 ka on benthic foraminifers from 152 to 157 cm. Several features in the record from GC1 suggest the presence of a discontinuity between 120 and 137 cm. The upper 120 cm of GC1 consists predominately of clay. Core X-radiographs show distinct burrows and occasional irregularly burrowed, lighter-colored layers that represent increased silt content (Osterman et al. 2006) (Fig. 23.4). The silt layers in this interval of the core are always thin (